Long-Term Subsidence and Compaction Rates: a New Model for the Michoud Area, South Louisiana Clint H

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LSU Master's Theses Graduate School

2008 Long-term subsidence and compaction rates: a new model for the Michoud area, south Louisiana Clint H. Edrington Louisiana State University and Agricultural and Mechanical College, [email protected]

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Recommended Citation Edrington, Clint H., "Long-term subsidence and compaction rates: a new model for the Michoud area, south Louisiana" (2008). LSU Master's Theses. 356. https://digitalcommons.lsu.edu/gradschool_theses/356

This Thesis 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 Master's Theses by an authorized graduate school editor of LSU Digital Commons. For more information, please contact [email protected]. LONG-TERM SUBSIDENCE AND COMPACTION RATES: A NEW MODEL FOR THE MICHOUD AREA, SOUTH LOUISIANA

A Thesis

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 Master of Science

The Department of Geology

By Clint H. Edrington B.S., University of New Orleans, 2005 May, 2008 ACKNOWLEDGEMENTS

I would like to thank my advisor, Dr. Michael “Mike” Blum, for the many resources provided during my study, particularly for his influence in my choice of a thesis topic, which is

such a critical issue facing south Louisiana, as well as the many stimulating ideas provided along

the way. I would also like to thank committee members Dr. Jeffrey Nunn and Dr. Jeffrey Hanor

for their advice and guidance, as they provided much insight from their respective areas of

research. Of course, many others contributed to this thesis in many ways, both large and small:

I would like to thank Dominion Exploration and Production for micropaleontological data,

GEOMAP Company for structure maps, both the Louisiana Geological Survey and the Louisiana

Department of Natural Resources for various subsurface data and guidance, both the Department

of Geology and Geophysics at Louisiana State University and the National Science Foundation

for financial support, and both faculty and fellow graduate students here in the Department, most

notably Dr. Blum’s students, i.e. my good friends. Finally, and certainly not least, I would like

to thank my family and friends who provided encouragement, support, and inspiration

throughout my study, and an outlet when I needed to break away from it.

TABLE OF CONTENTS

ACKNOWLEDGMENTS ………………………………………………………………………. ii

LIST OF TABLES ……………………………………………………………………………..... v

LIST OF FIGURES …………………………………………………………………………….. vi

ABSTRACT …………………………………………………………………………………… vii

INTRODUCTION ……………………………………………………………………………..... 1

BACKGROUND ………………………………………………………………………………... 3 Study Area ………………………………………………………………………………. 3 Evolution of the Gulf of Mexico Basin ...………………………………………………... 3 Migration of Depocenters in South Louisiana and the Adjacent Continental Shelf: Miocene – Present ...……………………………………………………………………... 6 Salt Tectonic Controls: Emphasis on the Northern Gulf Margin ...……………………. 10 Compaction …………………………………………………………………………….. 14 Modern Subsidence …………………………………………………………………….. 15

METHODS …………………………………………………………………………………….. 22 Long-Term Subsidence Rate Calculations ……………………………………………... 22 Long-Term Compaction Rate Calculations ……………………………………………. 27

RESULTS ……………………………………………………………………………………… 37 Subsidence and Compaction Rates …………………………………………………….. 37 Long-Term Subsidence Rates ………………………………………………….. 37 Long-Term Compaction Rates ………………………………………………… 37 Structural Cross Section ……………………………………………………………….. 37 First-Order Interpretations ...... 40 Differential Movement of Faults and Subsidence Rates ……………………… 40 A Revised Structural Cross Section …………………………………………… 41 Activation Order of Faults …………………………………………………….. 46 Differential Movement of Faults and Compaction Rates ……………………... 46

DISCUSSION ………………………………………………………………………………….. 48 The Michoud Fault ……………………………………………………………………... 48 Rate Comparison: Dokka (2006) and This Study ……………………………………… 48 Subsidence Rates: The Last 100 kyrs ………………………………………………….. 50 The Last 4,000 yrs ……………………………………………………………… 51 The Baton Rouge-Tepetate Fault System as a Model for the Michoud Fault …………. 54

CONCLUSIONS ………………………………………………………………………………. 56

REFERENCES ………………………………………………………………………………… 59

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APPENDIX: SOURCES OF ERROR: ASSUMPTIONS AND METHODS …………………. 66

VITA …………………………………………………………………………………………… 70

LIST OF TABLES

1. Long-term subsidence and compaction rates for the Michoud study area ……..………. 38

LIST OF FIGURES

1. LandSat image of southeast Louisiana and the Michoud study area …………….……… 4

2. Late Pliocene – Quaternary glacio-eustasy and the northern Gulf margin ……………… 7

3. The Miocene section of south Louisiana and the adjacent continental shelf .…………… 8

4. Locations of the depocenter during the Pliocene – Pleistocene …………………………. 9

5. The Cenozoic northern GOM margin divided into four structural provinces …………. 11

6. Evolution of allochthonous salt within the Eastern Province ………………………….. 12

7. Mechanisms, both anthropogenic and natural, contributing to modern land subsidence in south Louisiana ……………………………………………………………………… 16

8. Holocene RSL curve for the Mississippi Delta region ………………………………… 17

9. Subsidence model from Dokka (2006) for the Michoud area ………………………….. 20

10. LandSat images of the Michoud study area ……………………………………………. 23

11. Structure maps of the study area ……………………………………………………….. 28

12. Schematic representation of “decompacting” a sediment column ……………………... 31

13. Porosity versus depth profiles for south Louisiana …………………………………….. 33

14. Semi-log graph representations of the Dickinson (1953) porosity versus depth curve ... 34

15. Image of a 700 ft section taken from well log 146563 ………………………………… 35

16. Preliminary cross section of the Michoud study area ………………………………….. 39

17. Modification of Figure 11 (A) …………………………………………………………. 42

18. Final cross section of the Michoud study area …………………………………………. 44

19. Fault interpretations projected onto subsidence rates of the three Miocene horizons …. 45

20. Stratigraphic cross section of East New Orleans ………………………………………. 52

21. The New Orleans Barrier Complex ……………………………………………………. 53

22. Fault displacement history of the Baton Rouge-Tepetate Fault System ………………. 55

ABSTRACT

This study examined the stratigraphic record of the Michoud area in East New Orleans,

Louisiana to address questions concerning the magnitude of, and processes leading to land-

surface subsidence. The hypotheses under review are based on recent geodetic studies, which

challenge the widely held position that modern subsidence is primarily a function of shallow

sediment compaction. Testing these hypotheses involved constructing a structural cross section

of the Michoud area using well logs, chronostratigraphic data, and fault picks, so as to evaluate

differential motion along specific faults through time. Employing ages and corrected depths for

three key subsurface horizons, long-term (Middle Miocene to Present) time-averaged subsidence

rates were calculated: rates range from -0.140 to -0.177 m/kyrs (-0.140 to -0.177 mm/yr). Long- term subsidence rates are incompatible with those derived from geodetic studies: geodetically derived subsidence rates (-14.2 to -23.0 mm/yr) for the Michoud area are two orders of magnitude greater than long-term subsidence rates. Considering the scale of resolution of respective techniques, caution is advised when comparing respective subsidence rates.

Nevertheless, the new subsurface, structural model for the Michoud area suggests reactivation of local faults, including any recent movement of the Michoud Fault, is a transient phenomenon that is likely related to rapid Quaternary sediment loading. It is reasonable to compare mean long- term compaction rates, which is a component of total subsidence, derived from this research to geodetically derived compaction rates of pre-Holocene strata. Using a standard decompaction technique, mean long-term compaction rates for strata residing above the Middle Miocene

Bigenerina Humblei horizon were calculated: rates range from -0.0704 to -0.0914 m/kyrs

(-0.0704 to -0.0914 mm/yr), which are two orders of magnitude less than geodetically derived, pre-Holocene strata compaction rates (-4.6 mm/yr). The findings of this research, particularly the discrepancy between mean long-term compaction rates derived in this study and pre-

vii

Holocene compaction rates derived geodetically, raises questions into the interpretations and/or accuracy of the geodetic data for the Michoud area, and therefore, the subsidence rates determined from such data.

viii

INTRODUCTION

Land-surface subsidence is the cumulative effect of processes that operate over different spatial and temporal scales, and at different depths. At its essence, subsidence creates accommodation. Accommodation leads to sediment accumulation, and over a large spatial and temporal scale, subsidence ultimately forms a sedimentary basin. The structural evolution of the

Gulf of Mexico (GOM) basin, in particular, has been, and is currently affected by a number of subsidence mechanisms (Kulp, 2000). The scope of this thesis examines two such mechanisms in south Louisiana: faulting and compaction.

Recent research on subsidence mechanisms has produced seemingly incompatible results.

One view, based on studies of Holocene relative sea-level (RSL) change, contends the

Pleistocene – Holocene contact in south Louisiana, which represents the stratigraphic top of the pre-Holocene deltaic depocenter, is stable (Törnqvist et al., 2006; González and Törnqvist,

2006). Thus, compaction of overlying Holocene sediment is considered the primary mechanism driving land-surface subsidence. Holocene sediment compaction rates are generally modeled from -0.69 to -2.2 mm/yr (Meckel et al., 2006; 2007).

A second view, based on late 20th century geodetic data, acknowledges that compaction of Holocene sediment occurs, but argues it is a secondary mechanism superimposed upon a more significant, deep-seated tectonic component (Dokka, 2006; Dokka et al., 2006). For instance,

Dokka (2006) inferred contributions to subsidence of >-7 mm/yr from reactivation of growth faults and -4.6 mm/yr from compaction of pre-Holocene strata for the period 1969 – 1977 from the Michoud area of East New Orleans, Louisiana. These relatively high subsidence rates are supported by additional studies (see Shinkle and Dokka, 2004; Dixon et al., 2006), which report historical subsidence rates of -3 to -30 mm/yr from different parts of the delta region, including rates as high as -29 mm/yr for the greater New Orleans area.

These opposing views report subsidence rates that differ by at least one-order of magnitude, and are, on first glance, not reconcilable. Differences cannot be explained as a consequence of different study site locations, because the Törnqvist et al. (2004) and Dokka

(2006) studies, which are ~75 km from each other, are both on the downdip side of faults inferred to be active by Dokka et al. (2006). Different methodologies, with inherently different levels of time-averaging and temporal resolution, may partly explain the different subsidence rates and, consequently, interpretations. However, both approaches should be tested with independently-derived datasets. One important set of data consists of the longer-term stratigraphic record.

Research here examines the stratigraphic record of the Michoud area to address the following two questions. First, if deep-seated faults in south Louisiana crop out and influence modern subsidence, as advocated by Dokka (2006) and Dokka et al. (2006), does the stratigraphic record support such high rates, or is this a recent and perhaps strictly transient phenomenon? Second, are the relatively high compaction rates of pre-Holocene strata advocated by Dokka (2006) supported by the stratigraphic record? To address these questions, this study will: (a) develop a subsurface, structural model for the Michoud area, where much of Dokka’s

(2006) data and interpretations are derived; (b) use dated stratigraphic horizons to derive long- term (Middle Miocene – Present) subsidence rates and their relationship to proposed faults; and

BACKGROUND

Study Area

The Michoud area of East New Orleans, Louisiana lies within the modern Mississippi

Delta Plain, specifically within the St. Bernard Delta Complex (Frazier, 1967; Roberts, 1997), on

the northern GOM margin. Geographically, it is surrounded by Lake Pontchartrain to the north

and west, Lake Borgne to the east, and the Mississippi River to the south (Fig. 1). The open

GOM is ~83 km due south. According to Shinkle and Dokka (2004), the Michoud area is

characterized by some of the highest recorded rates of modern land-surface subsidence in the

United States. Its precarious setting, previous work by Dokka (2006) and Dokka et al. (2006),

and the long-term post-Katrina plans to rebuild the areas surrounding New Orleans, make this

site an appropriate focus for this study.

Evolution of the Gulf of Mexico Basin

A short discussion of the long-term evolution of the GOM provides context for the focus

on subsidence processes and rates. The GOM is a classic rift-to-drift passive margin basin.

Early stages of GOM basin evolution are traced to the Late Triassic – Middle Jurassic, when

mantle upwelling resulted in rifting of continental crust, extension and crustal attenuation, and

development of transitional crust (Worrall and Snelson, 1989; Buffler, 1991; Salvador, 1991).

Late Middle Jurassic marine incursions produced thick salt deposits in the subsiding rift basin,

followed by sea-floor spreading and emplacement of Late Jurassic oceanic crust during the early

drift phase evolution of the nascent passive margin (Worrall and Snelson, 1989; Buffler, 1991;

Salvador, 1991). During the Cretaceous, the GOM margin evolved into a carbonate shelf and

platform environment, linked to the Western Interior Seaway. Carbonate deposition ceased by

the end of the Mesozoic, and the Cenozoic has been dominated by siliciclastic sediment input

4 and progradation of the passive continental margin (Buffler, 1991; Salvador, 1991; Galloway et al., 2000; Galloway, 2001).

Galloway (2005) partitioned the Cenozoic GOM into distinct phases that reflect evolution of hinterland drainage, funneling of major river systems into deep-seated structural lows, and subsequent depositional history. Through the Paleocene – Oligocene, major depocenters were located in East Texas (the Houston Embayment) and South Texas (the Rio Grande Embayment), and progradation was concentrated on the western and northwestern shelf margins (Winker,

1982; Galloway et al., 2000, Galloway, 2005). During the Miocene, the primary axis for fluvial sediment input shifted to the Mississippi embayment, a position that persists today (Winker,

1982; Galloway et al., 2000, Galloway, 2005). The final Cenozoic phase was marked by

Pliocene tectonic rejuvenation of the Rockies and Plio-Pleistocene glaciation in the mid- continent, which resulted in rejuvenation of sediment sources and development of the present- day integrated Mississippi drainage as a continental-scale system (Winker, 1982; Galloway et al.,

2000, Galloway, 2005).

Continental glaciation and associated glacio-eustasy were the most important forcing mechanisms for the Late Pliocene – Quaternary Lower Mississippi Valley and northern GOM shelf and slope stratigraphy (Fisk, 1944; Suter and Berryhill, 1985; Coleman and Roberts, 1988;

Autin et al., 1991; Saucier, 1994; Blum and Törnqvist, 2000). In earlier literature, eight major eustatic cycles were documented in the Gulf and correlated to the North American mid-continent glacial-interglacial record (Beard et al., 1982; Saucier, 1994). Although the record of glaciation and global sea-level change is now known to be more complex (see Miller et al., 2005), more recent research has yet to refine the pre-Late Pleistocene record of the Mississippi system, including the offshore shelf and deep basin record. However, the present full interglacial condition is not representative for this time frame: the majority of the Pliocene – Late

Pleistocene was characterized by lower sea-level, stream extension across the shelf, and

progradation of the shelf margin (Fig. 2) (Suter and Berryhill, 1985; Blum and Törnqvist, 2000).

Migration of Depocenters in South Louisiana and the Adjacent Continental Shelf: Miocene – Present

Miocene sediments in south Louisiana and the adjacent shelf generally accumulated by

deposition in deltaic and shallow marine environments associated with the ancestral pre-glacial

Mississippi system (Rainwater, 1964; Winker; 1982; Galloway et al., 2000; Galloway, 2001).

The Miocene interval is composed almost entirely of sand, silt, and clay, reaches thicknesses as great as 6,100 m (Rainwater, 1964), and represents progradation of the shelf margin by as much as 200 km (Galloway, 2000). The Lower Miocene depocenter was located in southwestern

Louisiana and the adjacent shelf, whereas the significantly thicker Upper Miocene depocenter was to the east (Rainwater, 1964; Winker, 1982; Wu and Galloway, 2002) (Fig. 3). Miocene strata in south Louisiana record numerous transgressive-regressive cycles, with major transgressions recorded in subsurface data as regional shale horizons.

Pliocene – Pleistocene deltaic and shallow marine deposition resulted in progradation of

the shelf margin by an additional 80 km (Woodbury et al., 1973; Galloway et al., 2000). The

Early – Middle Pliocene depocenter was located offshore east-central Louisiana, with a total

thickness of 3,600 m, whereas the Late Pliocene – Pleistocene depocenter occurred farther west,

with a total thickness of 3,600 m as well (Woodbury et al., 1973; Galloway et al., 2000) (Fig. 4).

As noted above, the Late Pliocene – Quaternary time period was dominated by continental glaciation and associated glacio-eustasy, which impacted the Mississippi drainage basin

significantly, as well as the positions of sea level to which the Mississippi channel graded.

Falling stage-to-lowstand deposits accumulated as the Mississippi channel extended to shorelines

at the shelf margin or farther basinward positions, and are typically inferred to be represented by

predominately braided stream fluvial deposits (McFarlan and LeRoy, 1988; Saucier, 1994). By 6

9 contrast, deposits representing sea-level rise and highstand within the study area are typically inferred to signify inner shelf deltaic or shallow marine depositional environments (McFarland and LeRoy, 1988; Saucier, 1994).

Salt Tectonic Controls: Emphasis on the Northern Gulf Margin

The autochthonous Upper Jurassic Louann Salt has played an active role in the structural evolution of the northern GOM basin, particularly in the Cenozoic. Gravity and differential sedimentary loading are the primary forces that control deformation within salt zones; hence, the location of depocenters is the primary control on the location of salt-tectonic activity (Peel et al.,

1995; McBride, 1998). Both autochthonous and allochthonous salt progressively evacuated under loading, thus providing accommodation as depocenters migrated basinward (Peel et al.,

1995; McBride, 1998). Peel et al. (1995) divided the north and northwest Gulf margin into four provinces that share similar stratigraphic and structural characteristics (Fig. 5). The Michoud study area is located in the Eastern Province, which experienced Miocene – Early Pliocene extension that accompanied basinward salt movement.

Figure 6 provides a series of north-south cross sections through the Eastern Province, illustrating the coevolution of major structures, salt, and major depositional units (from McBride,

1998). Worral and Snelson (1989) suggested that most salt structures on the northern Gulf margin develop by a “down-building” mechanism as opposed to diapirism, as previously inferred. In their view, flat-topped salt massifs, similar to those present on the middle to upper slope today, begin to collapse by salt evacuation at the onset of significant depositional loading.

In conjunction with peripheral growth faults, this massif will continue to deflate, eventually evolving into a salt dome or weld structure that is flanked by a sediment trapping mini-basin

(Worral and Snelson, 1989).

As the depocenter migrates basinward, coarser shelf sediments prograde over the older slope facies (refer to Fig. 6 h-j), and remaining salt structures and growth faults are buried (Fails,

1990). Where down-building of salt structures continues, growth faults that developed initially on the upper-slope remain active and propagate upwards through new strata (Fails, 1990;

McBride, 1998), and younger faults associated with delta front instabilities form on the shelf margin (Jackson and Galloway, 1984; Galloway, 1986). Younger faults may link with buried growth faults, and, in doing so, keep ancestral faults active (Fails, 1990). With continued basinward migration of the depocenter, faulting generally ceases as sediment accumulation in updip positions becomes less significant (Fails, 1990).

Figure 6. Evolution of allochthonous salt within the Eastern Province. A – J shows the extrusion of allochthonous salt within the Eastern Province as four major episodes, with each episode succeeded by salt evacuation. Evacuation of the Terrebonne Salt Sheet provided accommodation beneath south Louisiana during the Miocene as major sedimentation shifted to the northern Gulf margin (6B – 6E). By the end of the Miocene (E – G), most of the salt within this allochthonous salt sheet had evacuated, and the deep structural framework observed today beneath south Louisiana was established. This framework was, subsequently, buried by continental shelf and continental sequences (6H – 6J). Star approximates the Michoud study area. Refer to Fig. 5 for location. Modified from McBride (1998).

Compaction

Overburden pressure is collectively supported by pore fluid pressure and the grain-grain bearing strength, or effective stress, of the sediment themselves (Terzaghi, 1948). As sediment is buried to increasing depths, increases in overburden pressure reduce porosity and total sediment volume through compaction (Athy, 1930). Therefore, an increase in compaction is characterized by an increase in effective stress and a decrease in pore pressure support as porosity is reduced.

For each value of porosity there is a corresponding maximum effective stress that a given sediment body can support without further compaction (Rubey and Hubbert, 1959). Thus, porosity is a surrogate measure for the degree of compaction in normally pressured strata.

Water-saturated mud is exceptionally porous at deposition (Rubey and Hubbert, 1959;

Dickinson, 1953; Baldwin and Butler, 1985; Kuecher, 1994). Mud-rich strata respond to initial loading by an increase in pore pressure, followed by expulsion of pore fluids until pore pressure reaches a corresponding lower limit for a given depth (Rubey and Hubbert, 1959). The lower limit that pore pressure reaches under normal circumstances is termed hydrostatic pressure

(Rubey and Hubbert, 1959). If dewatering is inhibited by an impermeable stratum, then overpressure develops as the retained fluid assumes a greater portion of the overburden pressure

(Hart et al., 1995). Overpressure, or compaction disequilibrium, is generally the result of sedimentation rates that exceed the ability of mud-rich sediments to expel pore fluids, and is common in GOM strata (Dickinson, 1953; Bethke, 1986; Harrison and Summa, 1991; Hart et al.,

1995). Two other mechanisms, although relatively unimportant in shallower depths, are smectite dehydration and aquathermal pressuring (Bethke, 1986; Harrison and Summa, 1991; Lahann,

2002).

Compaction of sands occurs by mechanical and chemical processes (Lundegard, 1992;

Gluyas and Cade, 1997). Initially, compaction of sands, like other sediment, is primarily

mechanical, from grain rearrangement (Gluyas and Cade, 1997), and burial will continue to

reduce intergranular volume by mechanical means until: (1) grains reach their absolute physical

lower limit, which is ~26% porosity in the absence of cement and matrix; or (2) mechanical

compaction is arrested before reaching its absolute lower limit by precipitation of pore-filling

cement (Paxton et al., 2002). Once mechanical compaction reaches its lower limit, further

reductions in volume occur through pressure solution, the dissolution of quartz along grain-to-

grain contacts (Paxton et al., 2002).

Modern Subsidence

Land-surface subsidence is the cumulative effect of processes that operate over different

spatial and temporal scales, and at different depths. Kulp (2000) summarized the range of

natural and anthropogenic mechanisms that contribute to modern subsidence onshore and

offshore Louisiana (Fig. 7). Shallow processes are dominated by compaction of Holocene

sediments and delta-front instabilities, whereas deep-seated processes include faulting, compaction of pre-Holocene sediments, fluid withdrawal, isostatic adjustments to sediment and water loading and the melting of the Laurentide Ice Sheet, thermal contraction, and halokinetics.

The latter two are considered insignificant beneath the Michoud study area: lithospheric cooling produces subsidence rates of -0.0001 mm/yr (Kulp, 2000), and most of the Terrebonne allochthonous salt sheet evacuated by the Late Miocene (McBride, 1998). Moreover,

instabilities within the Holocene deltaic strata occur farther basinward, within the modern delta complex at the shelf margin.

Modern subsidence is contentiously debated, not only in the rates measured but also with respect to the principal driving mechanism. From this debate, two end members have arisen.

The first view approaches subsidence by measuring Holocene RSL rise in the Mississippi Delta, using subsurface basal peat as a sea-level indicator (refer to Törnqvist et al., 2004 for a detailed

description; also Törnqvist et al., 2006; González and Törnqvist, 2006). Their rationale is that peat-forming wetlands onlapped the highly consolidated Pleistocene basement, of which the delta lies upon, during Holocene transgression. Hence, dating basal peats produces a RSL curve, and because the effect of Holocene sediment compaction is eliminated from the equation, the

RSL curve records only the interaction between eustasy and tectonic subsidence.

Törnqvist et al. (2004; 2006) and González and Törnqvist (2006) acquired basal peat data spanning a large section of the Mississippi Delta (i.e. on both the east and west margins of the incised Mississippi paleo-valley). Resultant RSL curves were compatible, and the composite

Holocene RSL curve shows a smooth, asymptotic rise to modern sea level (Fig. 8): between ca.

8,000 – 7,000 yrs BP, ca. 7,000 – 3,000 yrs BP, and ca. 1,400 – 400 yrs BP, RSL rose ~3.5 mm/yr, ~1.5 mm/yr, and ~0.55 mm/yr, respectively. Note that the most recently measured rate of late Holocene RSL rise, ~0.55 mm/yr, is as much as two orders of magnitude less than geodetically measured land-surface subsidence (see Shinkle and Dokka, 2004; Dokka, 2006;

Dokka et al., 2006). By assuming eustatic sea-level change was negligible between ca. 1,400 –

400 yrs BP, González and Törnqvist (2006) suggested the late Holocene RSL curve predominately tracks the glacio-isostatic contribution to subsidence. Thus, this view contends the Pleistocene-Holocene contact in south Louisiana is relatively stable (i.e. faulting is insignificant), and compaction of Holocene sediment is the dominant mechanism driving modern land-surface subsidence.

Shallow compaction in south Louisiana varies with the thickness of Holocene sediments, which

in turn reflects the history of delta evolution (Kuecher et al., 1993; Kuecher, 1994; Roberts et al.,

1994). Roberts et al. (1994) demonstrated that subsidence rates generally increase both across

and down valley where topstratum deposits thicken over the entrenched Pleistocene alluvial

valley (see Fisk, 1944). Superimposed upon Holocene accommodation is the delta cycle, an inherent suite of processes described by Roberts et al. (1994) and Roberts (1997) as producing a

series of “deltas within deltas,” which offset and overlap on different spatial and temporal scales.

Accordingly, delta switching leads to the gradational stacking, both vertically and horizontally,

of different lithofacies, with argillaceous and organic rich types being more prone to compaction

(Kuecher et al., 1993; Kuecher, 1994).

Using a Monte Carlo modeling approach, Meckel et al. (2006; 2007) numerically modeled sediment-column thickness, lithofacies, and accumulation rates to test the above

observations. Accumulation rates are shown to strongly influence compaction rates (e.g. high accumulation rates correlate with high compaction rates), and although a theoretical range of compaction rates is possible for a given thickness, compaction rates generally increase with increasing thickness. Their data also support a direct relation between stratigraphic heterogeneity and compaction rates: high rates of compaction are favored by high proportions of compaction-prone lithofacies (e.g. peat) that are loaded by a high-density, permeable lithofacies

(e.g. bar sand), whereas low rates of compaction are associated with low proportions of compaction-prone lithofacies combined with high proportions of lithofacies that form good hydrologic seals (e.g. prodelta mud). Modeled present compaction rates are, generally, from

-0.69 mm/yr to -2.2 mm/yr, and rarely exceed -5 mm/yr. However, a considerable percentage of their modeled compaction rates are slightly less, but on the same order of magnitude, than observed radiometric subsidence rates (see Kulp, 2000). Thus, Meckel et al. (2006; 2007)

18 suggested processes other than Holocene sediment compaction may have contributed to

Holocene subsidence.

The second view challenges the paradigm that modern land-surface subsidence in south

Louisiana is predominantly controlled by shallow sediment compaction (Dokka, 2006; Dokka et al., 2006). Sediment compaction is unequivocal in the literature, and the above authors do not refute its contribution to modern subsidence. However, they contend recent geodetic surveys support a more critical, deep-seated tectonic component.

Dokka (2006) constructed a modern subsidence model for the Michoud area of East New

Orleans (Fig. 9), choosing the area for three reasons: (1) it exhibits some of the highest subsidence rates in the region; (2) multiple geodetic surveys have been conducted here over the last 50 years; (3) subsidence could be determined as a function of depth, because closely spaced benchmarks are attached to rods and wells that extend to varying depths. Using multiple epoch, first-order leveling data (refer to Shinkle and Dokka (2004) for a detailed method description), he computed vertical displacements for an array of benchmarks that span the proposed Michoud

Fault. It is noted that Dokka (2006) discerned the subsurface location of the Michoud Fault from two unlikely sources: (1) a series of mega-regional cross sections that describe allochthonous salt evolution across the northern Gulf margin (see Fig. 6); and (2) a regional tectonic map of the northern Gulf Coast with a scale of 1:1,000,000 (see Hickey and Sabate, 1972). A fault trace

(i.e. black line) drawn on a map of this scale is approximately 0.5 kilometer in thickness. It is not explicitly stated how Dokka (2006) projected this fault onto the surface.

The nature of benchmarks in the Michoud area allowed Dokka (2006) to assess three components of subsidence (refer to Fig. 9): (1) a deep component that contains no contributions from processes above 2,011 m depth (Middle Miocene and above strata); (2) an intermediate component that includes processes acting between 2,011 m and 178 m depths (Middle Miocene –

Upper Pleistocene strata); (3) a shallow component that contains processes acting at or above

178 m depth (Upper Pleistocene – Holocene). The deep component is determined based on the motions of one benchmark tied to what appears to be, though not explicitly stated, a waste disposal well that penetrates to a depth of 2,011 m. The intermediate component is determined based on the difference in motions between benchmarks tied to three water wells, which penetrate to depths between 170 – 178 m, and the one nearby deep waste disposal well. It is noted here that water wells are designed to extract water from the subsurface; however, Dokka

(2006) argues the effects of water withdrawal on subsidence in the Michoud area have been

relatively insignificant. The shallow component is based on the motions of many benchmarks,

which appear to be mostly at the surface or anchored in shallow, mostly Holocene, depths.

Measured land-surface subsidence rates for two survey epochs, 1969 – 1971 and 1971 –

1977, were -23.0 mm/yr and -14.2 mm/yr, respectively: hence, geodetically measured land-

surface subsidence rates are at least one order of magnitude greater than modeled, shallow

sediment compaction rates. During these two epochs, respectively, the deep component contributed -16.9 mm/yr and -7.1 mm/yr and was attributed to slip on the Michoud Fault; the

intermediate component was constant during both epochs at -4.6 mm/yr and was attributed to

intermediate-sediment compaction; the shallow component contributed -1.5 mm/yr and -2.5 mm/yr, respectively, both of which correlate with modeled Holocene compaction rates, and was attributed to shallow sediment compaction, drainage of organic soils, and groundwater withdrawal. Hence, a disproportionate percentage of observed subsidence (93% and 82%

during respective epochs) is attributed to mechanisms that operate beneath shallow Holocene strata. Thus, this view contends a deep-seated tectonic component is driving subsidence, with a

secondary, shallow process superimposed.

METHODS

Questions stated in the introduction were addressed using subsurface data from the

Michoud area of East New Orleans, Louisiana, on the north (updip) and south (downdip) sides of

the proposed faults of Dokka (2006; see also Dokka et al., 2006). Subsurface data were acquired from a variety of sources, including state government, industry, and the published literature.

Methods were subdivided into those that lead to calculation of long-term subsidence rates, versus

those that lead to calculation of long-term compaction rates as a contribution to total subsidence.

Long-Term Subsidence Rate Calculations

Part one was designed to calculate long-term time-averaged subsidence rates and

differential motions in a north-south transect across inferred faults in the Michoud study area.

Initial data were gathered from Strategic Online Natural Resources Information System

(SONRIS), a Web-accessible database maintained by the Louisiana Department of Natural

Resources (www.dnr.louisiana.gov) that allows for the retrieval of information through GIS

access to interactive maps and/or by querying, viewing, and downloading electronic documents.

A GIS map of the study area, including the location of existing wells, was obtained (Fig. 10 (A)),

and the two faults interpreted to be active by Dokka (2006) were projected onto this map. Next,

geophysical logs for each well within the study area were examined: each log was acquired in

analog form, because they had been scanned as image files into the SONRIS database. Hence,

primary log data were not available. Wells that begin logging close to the surface (from 300 –

100 ft or 91.44 – 30.48 m) were retained, whereas others were discarded.

Micropaleontological data were critical to development of a geochronological model for this study. Micropaleontological data were provided courtesy of Dominion Exploration and

Production of New Orleans, Louisiana. Five of the wells that log sufficiently close to the surface, so as to be useful for this study, share a common micropaleontological datum, the first

Figure 10. LandSat images of the Michoud study area. (A) Well locations are indicated with their serial number. Two outcropping faults proposed by Dokka (2006) are indicated as heavy black lines. (B) Red line depicts the north-south transect of five petroleum wells across the two proposed faults. Images taken and modified from SONRIS (www.dnr.louisiana.gov).

25 downhole occurrence of the benthic Foraminifera Bigenerina Humblei. Waterman (2005) placed the first downhole occurrence of Bigenerina Humblei in the Middle Miocene, ca.

12.85 Ma. These five wells were retained, whereas others were discarded (Fig. 10 (B)).

Additional chronostratigraphic controls include the tops of the Lower and Middle

Miocene, and both were identified in well 47277 by Bebout and Gutiérrez (1983). Ages for the tops of the Lower and Middle Miocene are ca. 16.4 and 11.2 Ma, respectively (Waterman, 2005).

In well log 47277, the top of the Middle Miocene is characterized by a pronounced shale marker and was easily identified and correlated with the other four well logs used in this study.

However, the top of the Lower Miocene is not associated with a clear shale or resistivity marker on the well logs, and therefore, this horizon could not be correlated over the distances separating these five wells. Nevertheless, Bebout and Gutiérrez (1983) picked this horizon in well 134936, which is not a part of the transect, but is located ~4.85 miles east, southeast of well 146563 (refer to Fig. 10 (A)): this marker was projected into the cross-section constructed for the Michoud study area in the proximal latitude of well 146563.

Subsurface depths of the three key horizons were corrected for mean sea level and interpreted water depth during the time period in question. Miller et al. (2005) provide a detailed reconstruction of eustatic sea level for the entire Cenozoic: sea-level positions were +20.5 m at

11.2 Ma, +1.5 m at 12.85 Ma, and -17.0 m at 16.4 Ma. According to Tipsword et al. (1966), the

Bigenerina Humblei represents a middle-shelf environment (-20 to -100 m water depth); the tops of the Lower and Middle Miocene were, likewise, interpreted to represent middle-shelf environments. Therefore, the stratigraphic horizons were corrected for an assumed intermediate water depth of -60 m. Ages and corrected subsurface depths were then used to calculate long- term time-averaged subsidence rates for each well in the north-to-south transect across the

Michoud study area.

A structural cross section for the study area was constructed from well logs,

chronostratigraphic data, and fault picks, so as to evaluate differential motion along specific

faults through time. Structure maps, which contain fault picks, were provided courtesy of

GEOMAP Company (Fig. 11). GEOMAP Company interpreted from well logs four down-to-

the-south faults that trend east – west across East New Orleans, Louisiana, with throws that vary

from 60 – 365 m (200 – 1,200 ft) at depths of interest to this study, from 2,440 – 3,350 m (8,000

– 11,000 ft). The two southernmost faults were identified in well logs used for this study,

whereas the other two occur farther north beyond the study area. Faults were projected in the subsurface to reflect the typical listric geometry, where dip angles decrease with increasing depth: according to Durham and Peeples (1956), faults in the Gulf Coast region generally dip

70° from the surface to 3,000 ft, decrease to 60° – 55° from 4,000 ft to 6000 ft, and then to 45° below depths of 8,000 ft. The two faults interpreted to crop out in the study area by Dokka

(2006) were not identified in the subsurface by GEOMAP Company, but were extrapolated onto the structural cross-section constructed for this thesis.

Long-Term Compaction Rate Calculations

Mean long-term compaction rates were calculated for strata that reside above the first downhole occurrence of the Bigenerina Humblei. Standard decompaction techniques sequentially remove stratigraphic units, and permit units below to decompact as they rise to shallower depths (Fig. 12) (Steckler and Watts, 1978; Sclater and Christie, 1980; Audet and

McConnel, 1994). When overlying sediment is removed, strata below expand by absorbing water (i.e. by increasing porosity), but the total volume of sediment grains within the remaining

sediment column remains constant (Steckler and Watts, 1978; Sclater and Christie, 1980; Audet

and McConnel, 1994). Thus, decompaction of a hydrostatically-pressured stratigraphic section

Figure 11. Structure maps of the study area were provided courtesy of, and modified from GEOMAP Company. The red line indicates the transect of five petroleum wells. (A) Structure map of two micropaleontological horizons spliced together: the Operculinoides and Cristellaria “I” horizons. (B) Structure map of three micropaleontological horizons spliced together: Marginulina Ascensionensis, Cibicides Opima (upper), and Cristellaria ‘I’ (lower). For both maps, wells 63803, 47277, and 34669 indicate fault displacement. Faults are shown as thick black lines.

involves the migration of sediment bodies upward along an exponential porosity versus depth curve (Steckler and Watts, 1978; Sclater and Christie, 1980; Audet and McConnel, 1994).

Athy (1930) empirically measured bulk rock density as a function of depth from the following equation,

-cz ρbd = ρ0 + (ρg – ρ0)(1 – e ), (1) where ρbd is the dry bulk density, ρ0 is a constant (ρbd at the surface), ρg is the grain density of the rock, c is a constant with dimensions in length-1, and z is depth. From equation (1) and the relationship between dry bulk density, porosity, and grain density,

ρbd = (1 – f)ρg, (2)

Rubey and Hubbert (1959) derived the following exponential porosity versus depth equation

used for this study’s decompaction model,

-cz f = f0e , (3) where f and f0 are porosity at depth and initial porosity, respectively.

Equation (3) leaves two constants to be determined: initial porosity f0 and the empirical

constant c. A value of 41% was used here as the original porosity for sand, drawn from

Kuecher’s (1994) examination of Holocene Mississippi Delta sediments and studies that

examined surficial sands elsewhere (see Pryor, 1973; Atkins and McBride 1992). Porosity of

muds at the surface were assumed to be 80% initially, likewise drawn from Kuecher’s (1994; see

also Dickinson, 1953) study of sediments proximal to the Louisiana coast. Constant c is unique

to individually distinct lithologies. The Atwater (1985) and Dickinson (1953) porosity versus

depth curves for south Louisiana were used to derive values of c for sandstone and shale,

respectively (Fig. 13). However, calculating c for shale is complicated by the rapid rate of

compaction above 610 m (Magara, 1971). Therefore, Dickinson’s curve was divided into four

separate depth intervals and transferred to semilog paper (Fig. 14). This in turn lead to a separate

original porosity value for each interval (refer to Fig. 14 (B)); after which, four different c values were calculated for the four different depth intervals. See appendix for further discussion.

The five well logs retained were printed off in their entirety. Well log depths are recorded in feet, so each log was broken down into 100 ft intervals (30.48 m) (Fig. 15). The only discrepancy was at the top (i.e. near the surface) and at the very bottom, because of the inconsistent depths at which the logs begin and to the irregular depths at which the Bigenerina

Humblei is picked, respectively. The top interval was designated as the first 400 ft (121.92 m)

for all wells, whereas bottom interval thickness varies.

Next, sand-shale ratios for each interval were established for each well log using their respective SP curves (refer to Fig. 15). The SP log determines gross lithology by its ability to differentiate permeable (i.e. sandstone) from impermeable zones (i.e. shale) (Asquith and

Krygowski, 2004). Using an Excel spreadsheet, initial porosity and c were calculated for each interval, based on each interval’s measured sand-shale ratio. Original thickness for each interval

was then calculated by decompacting the sediment column using the following equation (Sclater

and Christie, 1980),

-cz1 -cz2 -cz1’ -cz2’ z2’ – z1’ = z2 – z1 – [f0/c] * [e – e ] + [f0/c] * [e – e ], (4) where z1 is the depth of the top of a given sediment interval, z2 is the depth of the base of the

same given sediment interval, z1’ is the new depth of the top of the same given sediment interval

migrated to a shallower depth, and z2’ is the new depth of the base of the same given sediment interval migrated to a shallower depth. Total compaction above the first downhole occurrence of

Bigenerina Humblei was ascertained by subtracting each well’s present sediment column above

this marker from each well’s respective decompacted sediment column. Mean long-term

compaction rates were derived by dividing the difference by 12.85 million years.

RESULTS

Subsidence and Compaction Rates

Long-Term Subsidence Rates

Long-term, time-averaged subsidence rates for the Michoud study area were calculated for three subsurface horizons: the Bigenerina Humblei and the tops of the Lower and Middle

Miocene. The first downhole occurrence of the Bigenerina Humblei was picked in wells

146563, 35029, 34669, 47277, and 63803, at depths of 2,182.4 m (7,160 ft), 2,096.1 m (6,877 ft),

2,050.7 m (6,728 ft), 2,328.7 m (7,640 ft), and 2,301.2 m (7,550 ft), respectively. Calculations

indicate long-term subsidence rates are -0.165 m/kyrs, -0.159 m/kyrs, -0.155 m/kyrs, -0.177

m/kyrs, and -0.175 m/kyrs, respectively (refer to Table 1). The Middle Miocene top was picked

for wells in the same order as above, at depths of 1,768.8 m (5,800 ft), 1,740.4 m (5,710 ft),

1,728.2 m (5,670 ft), 1,685.5 m (5,530 ft), and 1,665.7 m (5,465 ft), respectively; long-term

subsidence rates are -0.154 m/kyrs, -0.152 m/kyrs, -0.151 m/kyrs, -0.147 m/kyrs, and -0.145

m/kyrs, respectively. The Lower Miocene top was picked in wells 47277 and 134936, at depths

of 2,371 m (7,780 ft) and 2,972 m (9,750 ft), respectively; long-term subsidence rates are -0.140

m/kyrs and -0.177 m/kyrs, respectively.

Long-Term Compaction Rates

Mean long-term compaction rates of the sediment column were calculated above the first

downhole occurrence of the Bigenerina Humblei. For wells 146563, 35029, 34669, 47277, and

63803, long-term compaction rates are -0.0742 m/kyrs, -0.0704 m/kyrs, -0.0813 m/kyrs, -0.0914

m/kyrs, and -0.0863 m/kyrs, respectively.

Structural Cross Section

Figure 16 is the preliminary structural cross section constructed through the Michoud

study area (refer to Fig. 10 (B) for a map view transect). Three subsurface horizons, the

Bigenerina Humblei and the tops of the Lower and Middle Miocene, are depicted. Although the

Lower Miocene top was picked in only two wells, with one pick projected into the cross-section,

this horizon is extended laterally, nonetheless. Faults indicated by GEOMAP Company, as well

as those advocated by Dokka (2006), are depicted. There is no evidence to suggest that Fault

“A” crops out, and there is no evidence to suggest the two faults advocated by Dokka (2006)

extend to great depths. Nevertheless, all faults are projected up to the surface and/or downbasin

for interpretation.

First-Order Interpretations

Differential Movement of Faults and Subsidence Rates

Unlike Faults “A” and “B,” which are verified to exist in the subsurface, neither of the

faults advocated by Dokka (2006) was used in the initial interpretations because they cannot be

clearly delineated in subsurface data. However, initial interpretations were used later to constrain possible scenarios for these two faults. Fault “A” was active, in a normal state, between the top of the Lower Miocene and the Bigenerina Humblei. Three factors support this

interpretation: on the hangingwall as opposed to the footwall, the top of the Lower Miocene is at

a much greater depth; on the hangingwall as opposed to the footwall, subsidence rates,

consequently, are higher when measured with respect to the top of the Lower Miocene; on the

hangingwall as opposed to the footwall, a much thicker stratigraphic section exists between the

Bigenerina Humblei horizon and the top of the Lower Miocene. Fault “A” ceased movement

prior to the Bigenerina Humblei event, as indicated by the consistent basinward dip of this

horizon on its hangingwall. The consistent basinward dip of the Middle Miocene top further

supports tectonic dormancy.

In contrast to Fault “A,” Fault “B” was inactive between the Bigenerina Humblei event and the top of the Lower Miocene. This is indicated by the, relatively, very thin stratigraphic

40 section between these two horizons on Fault “A’s” footwall. However, Fault “B” later became active at or subsequent to the Bigenerina Humblei event. Three factors support this interpretation: the Bigenerina Humblei horizon is located at greater depths on the footwall of

Fault “A,” as opposed to its hangingwall; subsidence rates of the Bigenerina Humblei horizon are, consequently, higher on the footwall of Fault “A,” as opposed to its hangingwall; a thicker stratigraphic section exists between the Bigenerina Humblei horizon and the top of the Middle

Miocene on the footwall of Fault “A,” as opposed to its hangingwall. Fault “B” ceased movement prior to the top of the Middle Miocene, as indicated by the consistent basinward dip of this horizon. Further, the consistent basinward dip of the Middle Miocene top and the relatively consistent subsidence rates of this horizon throughout the entire Michoud study area reflects the basinward shift in the depocenter and, consequently, tectonic quiescence.

A Revised Structural Cross Section

The Bigenerina Humblei horizon has subsided to a greater depth in well 47277 than in well 63803, both of which are in the same fault block with respect to Faults “A” and “B” (refer to Fig. 16). Considering Fault “B” was active ca. 12.85 Ma, one expects the Bigenerina Humblei horizon between wells 47277 and 63803 to dip down (up-basin) into Fault “B”; however, this horizon does the opposite. A plausible explanation is that GEOMAP Company’s previously interpreted Fault “B” is actually two separate faults: Faults “B” and “B1” (Fig. 17). This implies

Faults “B” and “B1” were contemporaneously active, with the Bigenerina Humblei horizon subsiding to a greater depth on Fault “B1’s” hangingwall. Based on these interpretations, Figure

18 is the final structural cross-section constructed for the Michoud study area.

Calculated long-term subsidence rates support interpreted periods of fault movement for

Faults “A,” “B,” and “B1” (Fig. 19). Fault “A” was interpreted to have been active between the top of the Lower Miocene and the Bigenerina Humblei time period: relatively high subsidence

Figure 17. Modification of Figure 11 (A). Previously interpreted Fault “B” is shown here as two separate faults: Faults “B” and “B1.” Structure map was provided courtesy of, and modified from GEOMAP Company.

rates in well 146563 on the downthrown side of Fault “A” supports this interpretation. Faults

“B” and “B1” were interpreted to have been active between the Bigenerina Humblei event and the top of the Middle Miocene: relatively high subsidence rates in wells 63803 and 47277 on the downthrown side of these respective faults supports this interpretation. Relatively consistent subsidence rates in all wells with respect to the top of the Middle Miocene supports the interpretation of post-Middle Miocene tectonic quiescence due to the basinward shift in the depocenter.

Activation Order of Faults

Thorsen (1963) described the age of faulting on the northern Gulf margin as becoming

progressively younger basinward. At first glance, the interpretation presented above seems in

opposition to this model. However, fault offset in wells 63803 and 47277 (refer to Figs. 17 and

18) indicate Faults “B” and “B1” were active during the Lower Miocene, but then abated prior

the Middle Miocene. Fault “A” later became active, or at least continued activity into the Middle

Miocene. Thus, activation order of Faults “A,” “B,” and “B1” is actually chronologically

consistent with the growth fault character of the northern Gulf margin. For reasons unknown,

Faults “B” and “B1” reactivated post-Bigenerina Humblei event, while Fault “A” remained

dormant. As Thorsen (1963) suggested, there may be considerable overlap in the time of

movement between proximal, successive faults, which may explain the fault history of this area.

Differential Movement of Faults and Compaction Rates

The sediment column above the Bigenerina Humblei horizon demonstrates greater rates

of compaction on Fault “A’s” footwall, versus its hangingwall (refer to Table 1 and Fig. 18). As

discussed above, Meckel et al. (2006; 2007) showed that shallow sediment compaction rates

generally increase with increasing stratigraphic thickness. It is suggested here that this fundamental principle applies to long-term compaction rates. Faults “B” and “B1” were active between the Bigenerina Humblei event and the top of the Middle Miocene, while Fault “A” was inactive. By the end of the Middle Miocene, faulting had ceased throughout the study area.

Hence, the stratigraphic section between the Bigenerina Humblei horizon and the surface in the

downthrown area of Faults “B” and “B1,” of which wells 47277 and 63803 penetrate, is thicker

than elsewhere in the study area. Therefore, wells 47277 and 63803 should, and do maintain

higher compaction rates. Compaction rates on the hanging wall of Fault “A” are irregular (i.e.

compaction rates for well 34669 are higher than expected), which may reflect the inherent error

46 in the techniques employed. Nonetheless, these rates impart the approximate magnitude of compaction for the preserved stratigraphic section above the Bigenerina Humblei horizon, beneath the Michoud study area.

DISCUSSION

The Michoud Fault

If the Michoud Fault is deep seated, movement would be expected during the Lower –

Middle Miocene, as is characteristic of faults in this part of southeastern Louisiana (see Thorsen,

1963). Data below the Bigenerina Humblei horizon, in the vicinity of where the Michoud Fault

is projected, are insufficient to make that determination. In addition, GEOMAP Company does

not pick this fault in the subsurface. It is plausible, nonetheless, that the Michoud Fault was

active out in front of Fault “A” prior to the Bigenerina Humblei event, but has just not been

mapped in the subsurface. As described above, Fault “A” ceased activity prior to the Bigenerina

Humblei horizon; likewise, data suggest no activity on the Michoud Fault from the Bigenerina

Humblei horizon to the top of the Middle Miocene. Furthermore, the Michoud study area is tectonically dormant above the Middle Miocene top, as a result of the basinward shift in the depocenter. Therefore, recent movement on the Michoud Fault, as advocated by Dokka (2006) and Dokka et al. (2006), is considered geologically anomalous. These arguments equally apply to the “Unnamed” Fault from Dokka (2006), which was described only in Figure 1 of that article.

Rate Comparison: Dokka (2006) and This Study

Subsidence rates derived in this study are incompatible with those derived in Dokka

(2006): Dokka’s (2006) subsidence rates are two orders of magnitude greater than subsidence

rates derived in this research. Long-term subsidence rates, as determined with respect to the

Lower Miocene top (ca. 16.4 Ma), the Bigenerina Humblei event (ca. 12.85 Ma), and the Middle

Miocene top (ca. 11.2 Ma), were temporally variable: rates fluctuated from as high as -0.177

m/kyrs (-0.177 mm/yr) to as low as -0.140 m/kyrs (-0.140 mm/yr), a reflection of differential

movement of faults during the Middle Miocene. In contrast, Dokka (2006) advocated

subsidence rates in the Michoud study area as high as -23.0 mm/yr, which he predominantly attributed to recent movement on the Michoud Fault.

Shinkle and Dokka (2004) and Dokka (2006) explained how geodetic surveying

measures vertical changes on the scale of years to decades, and they advise against extrapolating

their rates beyond this scale. In contrast, techniques used in this study measured changes on the scale of tens of thousands to millions of years. The problem with comparing rates derived from different scales of resolution, such as comparing the subsidence rates from Dokka (2006) and

those from this research, is illuminated in Sadler’s (1999) study on sediment accumulation rates.

Sadler (1999) approached sediment accumulation from the perspective of fractal geometry, i.e.

as a process system with a common pattern nested within itself at different scales of resolution.

He described depositional hiatuses as permeating the stratigraphic record on all spatial and

temporal scales. As the temporal frame of measurement increases, larger spatial hiatuses tend to

be incorporated into calculations: consequently, short-term rates (i.e. mm/yr) are systematically

faster than long-term rates (i.e. m/kyr) (Sadler, 1999). This same approach is applied to

subsidence, which implies that instantaneous geological events (e.g. a fault slip) “average out”

over the long term. In this regard, subsidence rates from Dokka (2006) and long-term

subsidence rates derived from this research should be compared with caution. Nevertheless, it is evident that subsidence rates advocated by Dokka (2006) for the Michoud area, which he contends is primarily controlled by faulting, are not representative of total subsidence from the

Miocene – Present.

It is reasonable to compare mean long-term compaction rates from this study with

Dokka’s (2006) pre-Holocene compaction rates. Accumulation rates (i.e. rates of change in

overburden pressure) within the Michoud study area have generally decreased since the Middle

Miocene, as the depocenter shifted basinward. However, accumulation rates deviated from this

general trend during the Quaternary, due to loading-unloading oscillations that accompanied

glacio-eustasy. Therefore, overburden pressure for a given stratigraphic interval within the

Michoud study area has not steadily increased with time, but has changed as accumulation rates

changed. Changes in compaction rates for a given stratigraphic interval should, correspondingly, mirror changes in overburden pressure. Though techniques used in this study are not integrative and, therefore, do not capture changes in compaction rates (i.e. changes in overburden pressure),

this study, nevertheless, decompacted the preserved sediment column above the Bigenerina

Humblei horizon and calculated the “preserved,” long-term compaction rate.

Dokka (2006) advocates compaction rates of -4.6 mm/yr for the pre-Holocene sediment

column (from depths of 170 to 2,011 m). In contrast, compaction rates from this study are, on

average, -0.0807 m/kyr (-0.0807 mm/yr) for the sediment column above the Middle Miocene

Bigenerina Humblei horizon (depth of horizon spans from approximately 2,050 to 2,330 m).

These rates are two orders of magnitude less than Dokka’s (2006) pre-Holocene compaction

rates. Barring the anthropogenic process of fluid withdraw, which Dokka (2006) ruled out as

occurring within the Michoud study area from depths 170 to 2,011 m, no geological mechanism

is suggested capable of causing compaction rates for a ~2,000 m sediment column to fluctuate in

time like the slip of a fault does in regards to total subsidence. Hence, pre-Holocene compaction

rates should not fluctuate over geologically short time scales.

Subsidence Rates: The Last 100 kyrs

The northern Gulf margin over the last 100 kyrs experienced progressive lowering of sea

level from marine oxygen isotope stage (MIS) 5c to MIS 2, followed by a rapid rise in sea level

during MIS 1 (Blum and Törnqvist, 2000). Major sediment deposition was, therefore, at the

shelf and shelf margin, whereas accommodation, and therefore sediment accumulation, in the

Michoud area was minimal until Holocene transgression. Hence, the Pleistocene Prairie surface

in the Michoud area was predominantly subaerially exposed (Kolb et al., 1975; Kolb and

Saucier, 1982; Saucier, 1994), only experiencing significant shallow marine and deltaic deposition, and therefore loading, since middle – late Holocene (see Roberts, 1997; Stapor and

Stone, 2004). Compaction rates of -4.6 mm/yr over the last 100 kyrs would imply the top of the

Pleistocene Prairie surface subsiding to ~460 m depth (i.e. the creation of ~460 m of accommodation). Such a scenario is known not to exist, as the top of the Pleistocene Prairie surface is at a subsurface altitude of -10 to -12 m in the Michoud area (Kolb et al, 1975; see Fig

20). In contrast, compaction rates of -0.0807 m/kyr would imply the top of the Pleistocene

Prairie surface subsiding to ~8 m depth, a scenario that is more agreeable with the subsurface record.

Dokka (2006) estimated modern faulting’s contribution to total subsidence in the

Michoud area from 1969 – 1977 to vary between -7.1 mm/yr and -16.9 mm/yr. Using a conservative value as a representative of these rates (here -10 mm/yr), such relatively high subsidence rates extrapolated over the last 100 kyrs would imply top of the Pleistocene Prairie surface subsiding to ~1,000 m depth. As described above, this is known not to be the case (refer to Fig. 20). However, considering again the Sadler (1999) approach, extrapolating fault induced subsidence rates over a large temporal scale (i.e. 100 kyrs) is not applicable. Nevertheless, it is evident that fault induced subsidence rates are not representative of the last 100 kyrs.

The Last 4,000 yrs

Increasing the scale of resolution to 4,000 yrs is perhaps more appropriate. Stapor and

Stone (2004) describe the geological evolution of the New Orleans Barrier Complex (NOBC), a

middle Holocene depositional feature (~4,000 yrs BP) lying 1 – 10 m beneath the greater New

Orleans area, including the Michoud area (Fig. 21). Middle – late Holocene sea-level elevation

on the northern Gulf margin is debated in the literature (see Blum et al., 2003; Törnqvist et al.,

2004; among others), and is beyond the scope of this study. Nevertheless, it is accepted that sea level at 4,000 yr BP was near modern level, +/- ~2 m. Hence, present-day subsurface altitude of the NOBC is roughly equivalent to its altitude at time of deposition, discrepancies due to any changes in eustatic sea level and the effects of subsidence. Applying pre-Holocene compaction rates of -4.6 mm/yr from Dokka (2006) over the last 4,000 yrs would imply ~18.5 m of subsidence; applying -10 mm/yr, a conservative representative, of fault induced subsidence rates

from Dokka (2006) would imply ~40 m of subsidence. Total subsidence from these two

mechanisms over the last 4,000 yrs would be ~58.5 m. Again, considering the Upper Pleistocene

– Holocene stratigraphy beneath the Michoud area (refer to Figs. 20 and 21), this is known not to

be the case. Therefore, it is apparent that the subsidence model advocated by Dokka (2006) for

the Michoud area is not representative of the last 4,000 yrs.

Figure 20 (B) shows a line representing the maximum upper limit to which the preserved

Pleistocene Prairie Formation can be restored to. One could entertain the idea that the irregularity of the actual Holocene-Pleistocene contact in East New Orleans is strictly due to

faulting: that the Pleistocene surface is a series of basinward dipping half grabens. This would

require an assumption that erosional processes have played no role in creating the irregular

Pleistocene surface, i.e. that the Pleistocene surface was a planar feature prior to faulting. It is

considered extremely unlikely herein that erosion by stream migration had no influence on the

Pleistocene surface prior to Holocene transgression. Nevertheless, entertaining the idea that faulting controls the irregularity of the actual Holocene-Pleistocene contact, fault induced subsidence rates of -10 mm/yr over the last 500 yrs could have created the ~5 m of maximum relief observed on the preserved Pleistocene surface.

The Baton Rouge-Tepetate Fault System as a Model for the Michoud Fault

Deep-seated faults in the Michoud study area probably reactivated, assuming they have done so as Dokka (2006) and Dokka et al. (2006) contend, in the same manner as the Baton

Rouge-Tepetate Fault System (Fig. 22). This system experienced rapid, syndepositional movement from Eocene – Oligocene; it then laid dormant throughout the Neogene before reactivating in the Pleistocene (Hanor, 1982). Rapid Quaternary sediment loading, in response to sea-level oscillations, is the probable cause of fault reactivation (Nunn, 1985). Estimated rates of faulting are 25.4 – 63.5 mm/kyr for the Baton Rouge-Tepetate Fault System over the last one

million years (Hanor, 1982). Using the Baton Rouge-Tepetate Fault System as a general model for the reactivation of faults on the delta’s northern perimeter, these rates suggest the ~-100 mm of subsidence Dokka (2006) attributed strictly to movement of the Michoud Fault from 1969 –

1977 is abnormally high, and probably a once-in-a-thousand-year, or less, slip.

CONCLUSIONS

This study examined the stratigraphic record of the Michoud area of East New Orleans,

Louisiana to address questions concerning subsidence rates, and mechanisms influencing subsidence rates. Data support syndepositional movement on previously mapped faults during the Early – Middle Miocene. During the same period, data are insufficient to make a determination for movement on the Michoud Fault, but assuming this fault is deep seated, movement is expected. By ca. 11.2 Ma (top of the Middle Miocene), the depocenter had shifted basinward of the Michoud area: all local faults ceased activity, which commenced a period of tectonic quiescence that lasted up until at least the Quaternary.

Recent studies suggest rapid sediment loading during the Quaternary reactivated faults in south Louisiana, and Dokka (2006) and Dokka et al. (2006) contend modern faulting is the dominant mechanism influencing subsidence in the region. Data generated in this research can neither refute nor support the latter position, for detecting recent fault movement is beyond the resolving power of the methods employed. Studies elsewhere do support this view: Gagliano et al. (2003; 2003) identified more than one-hundred outcropping faults in southeastern Louisiana, with 61% correlating to known deep-seated faults. Nevertheless, the stratigraphic record suggests recent movement measured on the Michoud Fault is a transient phenomenon.

Compaction rates derived in this research do not support the subsidence model advocated by Dokka (2006) for the Michoud study area. This does not necessarily refute his main point: that the Michoud Fault is the dominant mechanism influencing present-day subsidence rates in the Michoud area. However, the discrepancy in compaction rates raises questions about the interpretations and/or accuracy of the geodetic data for the Michoud area, and therefore, the subsidence rates determined from such data. If pre-Holocene compaction rates from Dokka

(2006) are real, groundwater withdrawal is a plausible explanation, as these rates were

56 determined from three benchmarks tied to three water wells. As Dokka (2006) noted, Dial

(1983) found that groundwater withdrawal in the Michoud area actually declined slightly from

1968 to 1982. However, Nunn (2003) showed that clay confining layers continue to lose pore water to adjacent aquifers, even when groundwater levels in the draining aquifers stabilize or rebound slightly. Such was the case beneath the Baton Rouge area: following a period of peak groundwater withdrawal for industrial use during the 1940s through 1960s, aquifer water levels leveled off or rebounded; however, subsidence was modeled to have continued at -5.0 mm/yr over the next two decades, until subsidence rates increased to near previous levels as industrial usage increased again (Nunn, 2003). Modeled subsidence rates (-5.0 mm/yr) due to compaction of the confining layers in the Baton Rouge area are remarkably similar to pre-Holocene compaction rates (-4.6 mm/yr) advocated by Dokka (2006) for the Michoud area.

Future studies should focus on the frequency of faulting in south Louisiana: i.e., what is the appropriate resolution scale for projecting recently observed, fault induced subsidence rates?

Based on Hanor’s (1982) study of the Baton Rouge-Tepetate Fault System, this study suggests recent movement on the Michoud Fault, from 1969 – 1977, is perhaps a once-in-a-thousand-year, or less, slip. Consequences are significant: one can basically ignore faulting if modeling projected, near-term subsidence rates for the Michoud study area. If, however, subsidence rates for the entire south Louisiana as calculated by Shinkle and Dokka (2004) are correct, which are generally one order of magnitude greater than Holocene sediment compaction rates, then ignoring faulting in models projecting subsidence rates would be imprudent: such high rates for the entire south Louisiana would imply a dominant tectonic component acting over the entire south Louisiana. It is regarded unlikely that separate fault systems within the entire south

Louisiana would reactivate in such uniform manner, and then, in concert, lie dormant for the next one thousand plus years. Hence, if rates advocated by Shinkle and Dokka (2004) are correct,

57 modern faulting events may be a more frequent process. More rigorous research (e.g. a detailed shallow-seismic study of the Michoud Fault) is needed, therefore, to more adequately quantify the frequency, in both a temporal and spatial sense, of Late Quaternary faulting in south

Louisiana.

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APPENDIX

SOURCES OF ERROR: ASSUMPTIONS AND METHODS

Horizons used for calculations, and the lithology described for each well are considered reasonably robust, but calculations are subject to error from the assumptions made and the methods employed.

Assumptions

Sediment grain size is either sand or clay. This form of simplicity is used in many basin

analysis models (see Steckler and Watts, 1978; Sclater and Christie, 1980; Audet and

McConnel, 1994).

Intergranular volume between framework grains is the sum of pore space, cements, and

depositional matrix (Paxton et al., 2002). For this study, intergranular volume was

assumed to equal porosity: (a) a detailed petrography study was beyond the scope of this

thesis, and sands were instead treated as texturally homogenous, rigid, spherical quartz

grains; (b) strata were treated as an open system that removes dissolved quartz from the

system and keeps pore fluids below saturation. Simplifying assumptions of this kind are

common in the literature (Gluyas and Cade, 1997; Paxton et al., 2002): for example,

cementation is routinely ignored to simplify porosity-depth models for compacting sands;

however, it is recognized that compactional processes play the dominant role in porosity

loss, whereas cementation plays a secondary role (Lundegard, 1992). Moreover,

sandstones from basins with low geothermal gradients, such as the GOM passive margin,

should experience low rates of pressure solution, because chemical reactions generally

are slower at lower temperatures (Lundegard, 1992; Burrus, 1998). As noted by

Bjørlykke and Egeberg (1993), quartz cementation in the Gulf Coast basin is poorly

developed down to depths of 2.5 – 3.0 km.

Porosity loss is a function of mechanical compaction only. See above.

The stratigraphic section above the Bigenerina Humblei horizon is hydrostatically

pressured. This assumption is in agreement with the Dickinson (1953) study, which

shows abnormal conditions in the vicinity of the Michoud study area occurring at depths

below ~2,900 m (9,500 ft). In addition, there is no indication of abnormal pressure from

well log headers.

The Bigenerina Humblei horizon, the top of the Middle Miocene horizon, and the top of

the Lower Miocene horizon were deposited in the following relation to modern sea-level,

respectivley: -58.5 m; -39.5 m; -77 m.

Because of well 63803’s SP curve’s poor quality above 792 m (2,600 ft) depth, data for

this depth interval were taken from the closest well, well 47277, and used in its place.

Methods

Equation (3) was originally derived from studies of compacting shale (see Athy, 1930;

Ruby and Hubbert, 1959). However, Sclater and Christie (1980) argued this exponential

equation is generally applicable to compacting sandstone. As Bahr et al. (2001)

explained, a simple fit curve, like equation (3), does not reproduce every detail in the

porosity profile, but it will capture the general trend. An exponential curve for describing

compacting sandstone has been advocated elsewhere (see Baldwin and Butler, 1985;

Gluyas and Cade, 1997; Paxton et al., 2002).

Figure 14 is an empirical relationship describing shale compaction (for further discussion,

see Dickinson, 1953; Magara, 1971). After transferring Dickinson’s (1953) shale

compaction curve (refer to Fig. 13 (B)) to semi-log graphing paper (refer to Fig. 14 (A)),

a straight line was drawn over the normal compacting trend (i.e. from the top of abnormal

pressure conditions, which is ~3,048 m depth beneath the surrounding New Orleans area,

to 610 m depth). Because shale experiences rapid rates of compaction above 610 m

depth, this straight line normal compaction trend breaks down. However, if one projects

the normal compaction trend upward, it intersects the surface at f0 = 0.39. A similar

empirical technique was used above 610 m depth (refer to Fig. 14 (B)). Between depths

of 610 m and 305 m, f0 = 0.45; between depths of 305 m and 152 m, f0 = 0.53; between

depths of 152 m and 0 m, f0 = 0.80. As discussed above, the decompaction technique

involves moving sediment intervals up a porosity versus depth curve, in the case of this

-cz research f = f0e . Therefore, when sediment intervals were decompacted (i.e. when

being brought to the surface), f0 used in decompaction calculations were 0.39, 0.45, 0.53,

and 0.80 when depths were >610 m, between 610 m and 305 m, 305 m and 152 m, and

152 m and 0 m, respectively.

The averaged, mean long-term compaction rate calculated from using this method

was -0.0807 m/kyr. To quantify error in the above method, mean long-term compaction

rates were calculated using only one of the four f0 (i.e. calculations used only one f0 for decompaction at all depths). When setting shale f0 = 0.80 for all calculations, the

averaged, mean long-term compaction rate was -0.1472 m/kyr. When setting shale f0 =

0.53 for all calculations, the averaged, mean long-term compaction rate was -0.0614 m/kyr. When setting shale f0 = 0.45 for all calculations, the averaged, mean long-term

compaction rate was -0.0497 m/kyr. When setting shale f0 = 0.39 for all calculations, the

averaged, mean long-term compaction rate was -0.0427 m/kyr. Magnitude of rate

difference between the two members is 0.1046 m/kyr. Variance from -0.0807 m/kyr is not greater than a magnitude of 0.0665 m/kyr. Hence, calculated mean long-term compaction rates herein are considered relatively robust. Furthermore, for reasons

68 described above, using a composite of f0 for the appropriate depth interval during decompaction is considered the superior method.

VITA

Clint Edrington was born and raised on the Mississippi Gulf Coast, where he attended grade school and eventually graduated from Saint Stanislaus High School in 1996. Like eustatic sea level, his higher education can be described as a series of oscillations, but Clint ultimately earned a Bachelor of Science in geophysics from the University of New Orleans in 2005. Upon graduation, Clint entered the master’s program in the Department of Geology and Geophysics at

Louisiana State University, where he pursued an interest in sedimentology and processes related to. In January of 2008, Clint began his doctorate in the Department of Oceanography and

Coastal Sciences at Louisiana State University, where he will continue his interest in sedimentology, with the ultimate goal of applying learned concepts to help restore Louisiana’s, and elsewhere, coastal environments.