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Primary and tectonic fabric intensities in mudrocks
TECTONOPHYSICS ELSEVIER
Tectonophysics 247 (1995) 105-119
Primary and tectonic fabric intensities in mudrocks Scott R. Paterson a,., Hao Yu a, Gerhard Oertel b a Department of Earth Sciences, University of Southern California, Los Angeles, CA 90089-0740, USA b Department of Earth and Space Sciences, University of California, Los Angeles, Los Angeles, CA 90024-1567, USA Received 18 May 1994; revised version accepted 20 September 1994
Abstract
In this paper, we present a summary of quantitative measurements of the preferred orientation of phyllosilicate grains in a variety of tectonically weakly deformed to undeformed, fine-grained, pelitic rocks. We also briefly examine a small amount of data on clay fabrics in deep sea cores, and compare them to fabrics in deformed equivalents from the Mariposa Formation and Calaveras Complex, Sierra Nevada, California. The measured fabric ellipsoids support the following conclusions: (1) a bedding-parallel foliation forms in pelitic rocks during compaction that in our samples is associated with shortening between 3% and 74%; (2) the amount of compaction is in part controlled by the quartz-feldspar content: interbedded sandy layers undergo little to no compaction and do not develop bedding parallel foliations; and (3) it is impossible to determine precisely the amounts of compaction, horizontal shortening, and later tectonic strain in deformed samples because of the variability of initial compactions. Our data, and results presented by others, indicate that strain paths in pelitic rocks always include early compaction, followed by an episode of simultaneous compaction and plane strain in accretionary wedges, and by a variety of other types of strain in other settings.
1. Introduction
The shapes and preferred orientations of objects have been used extensively to evaluate strain in rocks. The preferred orientation of platy minerals has proven useful in this regard in finegrained sedimentary rocks, which often lack other suitable markers. However, there are a variety of processes that orient grains in sedimentary rocks including alignment during deposition, compaction, and pre- and post-lithifaction tectonic strain. Determining the magnitude of tectonic
* Corresponding author.
strain requires some information about prelithification fabrics. T h e r e exists a vast literature on the changes in microstructure and physical properties of muds and silts. For example, older summaries by Rieke and Chilingarian (1974) and Margara (1978) and recent work summarized in fifty-nine papers in Bennett et al. (1991) discuss the processes, microstructures, and resulting physical properties that occur during deposition, compaction, and lithification of fine-grained sediments. Baker et al. (1993) summarize changes in porosity, seismic velocity, seismic velocity anisotropy, and microstructure during compaction of pelites as well as new experimental results. Oertel (1983) has summarized techniques for measuring preferred orientations of phyllosili-
0040-1951/95/$09.50 © 1995 Elsevier Science B.V. All rights reserved SSDI 0 0 4 0 - 1 9 5 1 ( 9 4 ) 0 0 1 9 9 - 5
106
S.R. Paterson et al. / Tectonophysics 247 (1995) 105-119
cate grains and the validity of assumptions made in using the March model to relate clay fabrics to strain. These summaries and previous studies support the following: (1) compaction of pelitic rocks increases with depth during burial; (2) clay particles rotate towards parallelism with bedding during compaction; and (3) the magnitude of compaction is to a large degree controlled by the ratio of phyllosilicate minerals to more equidimensional grains (e.g., Curtis et al., 1980). Oertel (1983) and Baker et al. (1993) also concluded that the preferred orientation of phyllosilicates successfully reflects strain caused by compaction and tectonic processes. What is particularly useful to structural geologists, and much less commonly published, are quantitative measurements of the three-dimensional preferred orientation of phyllosilicate grains. Oertel and coworkers (i.e., Oertel and Curtis, 1972; Reed and Oertel, 1978; Curtis et al., 1980; Hildebrand-Mittlefehldt, 1980; Krishnan and Oertel, 1980; Engelder and Oertel, 1985; Evans et al., 1989; Oertel et al., 1989) have published some such data collected in rocks that at the scale of individual samples are not strongly deformed by tectonism. This paper presents a summary of data collected by Oertel and Paterson in a variety of tectonically weakly deformed to undeformed, fine-grained, pelitic rocks. We also examine a small amount of data on clay fabrics in deep sea cores, and compare them to fabrics in their deformed equivalents from the Mariposa Formation and Calaveras Complex, Sierra Nevada, California. Our results are compatible with the conclusions noted above. Our intention is not to reexamine these conclusions but, instead, provide a summary of quantitative measurements and to explore the implications of these measurements for studies of deformed equivalents.
2. Description of samples Deep Sea Core 174A: This core, drilled in the distal portion of the Astoria Fan off the coast of Oregon and seaward of the toe of the accre-
tionary prism, reached 879 m below the sea floor and recovered rocks from two units: (1) Upper Pleistocene thick- to thin-bedded, medium to fine turbiditic sands consisting of continental detritus deposited in the Astoria fan (0-284 m); and (2) the underlying Plio-Pleistocene thin-bedded sands, silts, and clays (284-879 m) representing abyssal plain deposits (Kulm et al., 1973). Ten samples were collected from the core, only four of which gave useful results (Table 1). During DSDP coring, an unknown amount of rotation around the vertical core axis occurred. Thus, orientations can only be measured with respect to horizontal. The preservation of horizontal bedding, laminations, and other sedimentary structures (Kulm et al., 1973) indicates that these samples were not significantly disrupted during coring or sample collection. Great Valley sequence: Six samples of claybearing siltstones were collected from a transect through the Cretaceous Great Valley sequence, California, in the Wilbur Springs Quadrangle. These interbedded sandstones and siltstones represent proximal and distal fore-arc fan sequences that after lithification have been tilted and locally deformed by bedding parallel faults (Ingersoll, 1979; Glen, 1990). However, at least in the sampled area, the rocks are not cleaved and show no microstructural evidence of crystal-plastic strain (Paterson and Yu, 1994). The sampled layers consist of 15-25% phyllosilicate grains, 50-75% quartz and feldspar grains, and small amounts of lithic fragments and volcanic detritus. Pigeon Point Formation: The Upper Cretaceous Pigeon Point Formation consists of turbiditic sequences of sandstone, siltstone, mudstone, and local conglomerate (Wentworth, 1960; Tyler, 1972). Facies F (see Howell and Joyce, 1981) is characterized by abundant slump structures. This facies consists of silt and sand distal turbidites with local, discontinuous mudstone and conglomerate layers. The mudstone consists of about 50% quartz, 10-20% feldspar, and variable amounts of micas, chlorite and clays. Organic material makes up a small but visible part in many samples, and some organic-rich layers are present. Although these rocks have been mildly deformed by open, upright, regional-scale folds
S.R. Paterson et aL / Tectonophysics 247 (1995) 105-119
and brittle faults (Howell and Joyce, 1981), metamorphic tectonites have not developed (Paterson and Tobisch, 1993), and the fabrics formed by penecontemporaneous slumping or compaction (e.g., Tobisch, 1984; Paterson and Tobisch, 1993). Eight samples have been analyzed from outcrops of facies F along a beach immediately south of Pigeon Point (Table 1). O D P H o l e 808c: To add to our data set of deep sea sediments we have included measured clay fabric orientatons from 9 samples collected by Morgan and Karig (1993) from ODP Hole 808c, drilled in the toe of the Nankai accretionary prism. Samples from depths ranging from 318 to 1137 m below the sea floor consist of hemipelagic silty-claystone with phyllosilicates making up 30% to 40%. They include both accreted rocks and subducted abyssal plain deposits. Behrmann and Kopf (1993) also measured clay fabrics in 11 samples from this site with samples obtained from depths of 23 to 533 m below the sea floor. Unfortunately, these authors did not present all three ratios of the calculated ellipsoids, thus precluding direct comparison of their results with data presented in this paper. Labrador Trough: These samples were collected in shaly beds of the Aphebian (early Proterozoic) exposed in the portion of the Labrador Trough that forms the Schefferville mining district of northeastern Canada (Krishnan and Oertel, 1980). They come mostly from the lowest stratigraphic unit in the trough, the Attikamagen Formation, some stratigraphic distance and at least one, or perhaps two, unconformities below the economically important Iron Formation. A few come from the Menihek Formation, which immediately overlies the Iron Formation. These calcareous, silty, and locally carbonaceous shales are deep water deposits formed in grabens in Archean basement. Whereas the central and eastern portions of the trough are moderately to strongly folded, its western rim is undeformed or only gently folded. The listed samples come mostly from this western margin of the trough or from long, straight, tectonically weakly strained limbs of moderate folds nearer the center of the trough. We excluded only those samples in which the shortest axis of the fabric ellipsoid (principal axis
107
of greatest contraction) departs significantly from the bedding-normal. This selection does not discriminate against some samples in which tectonic increments, even fairly large ones, happened to have been coaxial with the compaction strain. New York Plateau mudstones: Devonian mudstones were collected mostly from exposed, horizontal or gently dipping beds (maximum of 3°) of the Catskill Delta, New York Plateau, although a few came from borehole cores in the same region (Engelder and Oertel, 1985; Oertel et al., 1989; Evans et al., 1989). This delta is a large clastic wedge with both marine and fluvial sediments deposited on continental basement. Most show a small amount of measurable tectonic strain, caused by bedding-parallel shortening approximately 50-100 Ma after deposition (Engelder and Oertel, 1985). But this tectonic strain is only a small fraction of the strain caused by compaction during burial. Yorkshire: This group are shales and mudstones of Pennsylvanian (Westphalian A) age collected in tectonically undeformed horizontal beds in the vicinity of Penistone (near Sheffield), Yorkshire, England (Oertel and Curtis, 1972, and unpublished). These samples come from units deposited in shallow epicontinental basins on a Paleozoic basement. Two of the samples are seat earths (marked in Table 1 by an appended -S), the clay-rich substrata of coal seams. Green River Basin: Strain in this group was determined at various points in a single, large (0.6 x 0.4 x 0.3 m) sample (Hildebrand-Mittlefehldt, 1980) of gently east-dipping beds of the Middle Fork Tongue member of the Eocene Green River Formation, collected in the Uinta Basin near Utah State Highway 50. These thin shale layers, interbedded with marly limestones, are part of a thick sequence of lake beds. The measurements were performed by HildebrandMittlefehldt, using the same X-ray technique as for all the other fabric measurements discussed in this paper. The results, however, are unavailable as tabulated numerical values, and were, therefore, estimated from their graphic representation on Fig. 4 of the quoted paper (the scale is more than adequate for the required two significant figures of strain). Although the main objective of
S.R. Paterson et al. / Tectonophysics 247 (1995) 105-119
108 Table 1 Primary fabric data from mudstones Axial ratios
Apparent
extensions (%)
Y
Z
Apparent
ellipsoids
Volume
Sl
LP
for X = 1
X
Y
Z
X
OP1
1.88
1.61
1.00
29.97
11.30
- 30.87
0.47
0.51
- 54
OP6
1.13
1.10
1.00
5.10
2.31
- 6.99
0.09
0.56
- 14
OP7
1.02
1.01
1.00
0.99
0.00
- 0.99
0.01
0.00
OP10
1.71
1.24
1.00
33.10
GVS25
1.57
1.46
1.00
19.07
GVS27
1.46
1.30
1.00
17.92
GVS30
1.61
1.36
1.00
GVS32
1.54
1.36
1.00
GVS34
1.58
1.57
GVS36
1.31
P22 P23
DSDP Cores 174A
-3.48
-0.20
- 3
-22.16
0.38
-58
10.73
-24.16
0.34
0.68
-41
5.00
- 19.23
0.27
0.39
- 39
23.99
4.73
- 22.99
0.34
0.29
- 48
20.37
6.30
- 21.84
0.31
0.42
- 43
1.00
16.72
15.98
- 26.13
0.37
0.97
- 37
1.25
1.00
11.14
6.05
- 15.16
0.20
0.65
- 27
1.74
1.53
1.00
25.55
10.39
- 27.85
0.41
0.54
- 49
1.45
1.26
1.00
18.61
3.07
- 18.20
0.27
0.24
- 40
P24 P25
1.58 1.66
1.39 1.47
1.00 1.00
21.55 23.30
6.94 9.19
- 23.07 - 25.72
0.33 0.37
0.44 0.52
- 44 - 47
P26
1.20
1.18
1.00
6.86
5.08
- 10.95
0.14
0.82
- 18
OP6
1.13
1.10
1.00
5.10
2.31
-6.99
0.09
0.56
- 14
OP7
1.02
1.01
1.00
0.99
0.00
- 0.99
0.01
0.00
- 3
P21
1.88
1.61
1.00
29.97
11.30
- 30.87
0.47
0.51
- 54
Great Valley
Pigeon Point
ODP Hole 808C 1
1.27
1.19
1.00
10.49
3.69
- 12.71
0.17
0.46
- 26
2
1.33
1.19
1.00
14.39
1.79
- 14.11
0.20
0.19
- 33
3
1.53
1.40
1.00
18.80
8.60
- 22.50
0.32
0.58
- 40
4
1.27
1.09
1.00
14.27
- 2.52
10.22
0.17
- 0.32
- 33
5
1.35
1.24
1.00
13.67
4.55
- 15.85
0.22
0.44
- 32
6 7
1.23 1.32
1.11 1.24
1.00 1.00
10.57 12.34
0.24 4.94
-9.78 - 15.17
0.14 0.21
0.04 0.51
-26 - 29
8 9
1.23 1.37
1.21 1.28
1.00 1.00
7.81 13.71
6.11 6.21
-12.59 - 17.19
0.17 0.24
0.85 0.57
-20 - 32
3.13
3.11
1.00
47.00
46.00
- 53.00
0.93
0.99
- 68
2.92
2.92
1.00
43.00
43.00
- 51.00
0.87
1.00
- 66
3.24
3.17
1.00
49.00
46.00
-54.00
0.95
0.97
-70
3.20
3.20
1.00
47.00
47.00
- 54.00
0.95
1.00
- 69
2.98
2.55
1.00
52.00
30.00
- 49.00
0.84
0.71
- 71
3.55
3.30
1.00
56.00
45.00
- 56.00
1.00
0.88
- 74
2.90
2.59
1.00
48.00
32.00
- 49.00
0.83
(I.79
- 69
2.33
2.00
1.00
40.00
20.00
- 40.00
0.64
0.64
- 63
2.27 2.87 2.41
2.15 2.50 2.12
1.00 1.00 1.00
34.00 49.00 40.00
27.00 30.00 23.00
- 41.00 - 48.00 - 42.00
0.65 (I.81 0.67
0.87 0.74 0.71
- 58 - 70 - 64
2.67 2.26
2.37 1.85
1.00 1.00
44.00 40.00
28.00 15.00
- 46.00 - 38.0(I
0.76 0.60
0.76 0.52
- 67 - 64
4.27 2.16 2.43
3.39 2.03 1.90
1.00 1.00 1.00
75.00 32.00 46.00
39.00 24.00 14.00
- 59.00 - 39.00 - 40.00
1.10 0.61 0.65
0.68 0.84 0.44
- 81 - 57 - 68
Labrador Trough
loss
S.R. Paterson et al. / Tectonophysics 247 (1995) 105-119
109
Table 1 (continued) Axial ratios
Apparent extensions (%)
A p p a r e n t ellipsoids
V o l u m e loss
X
Y
Z
X
Y
Z
Sl
LP
for X = 1
2.02
1.88
1.00
29.00
20.00
- 36.00
0.55
0.79
- 54
2.11
1.98
1.00
31.00
23.00
- 38.00
0.59
0.83
- 56
2.35
1.95
1.00
41.00
17.00
- 40.00
0.64
0.56
- 65
2.16
1.94
1.00
34.00
20.00
-38.00
0.59
0.71
-59
New York Plateau
1.85
1.71
1.00
26.00
16.00
- 32.00
0.47
0.73
- 50
1.65
1.49
1.00
22.00
10.00
- 26.00
0.37
0.59
- 45
2.00
1.91
1.00
28.00
22.00
- 36.00
0.55
0.86
- 52
1.87
1.79
1.00
25.00
20.00
- 33.00
0.49
0.87
- 49
2.31 1.64
2.12 1.50
1.00 1.00
36.00 21.00
25.00 11.00
-41.00 - 26.00
0.65 0.37
0.80 0.65
-60 - 44
2.83
2.52
1.00
47.00
31.00
- 48.00
0.81
0.78
- 68
1.71 2.06
1.58 1.84
1.00 1.00
23.00 32.00
14.00 18.00
- 28.00 - 36.00
0.41 0.55
0.72 0.69
- 46 - 57 - 48
1.87
1.81
1.00
25.00
21.00
- 33.00
0.50
0.90
2.06
1.94
1.00
30.00
22.00
- 37.00
0.57
0.82
- 55
1.93
1.73
1.00
29.00
16.00
- 33.00
0.50
0.68
- 53
2.28
2.03
1.00
37.00
22.00
- 40.00
0.63
0.72
- 61
2.34
2.08
1.00
38.00
23.00
- 41.00
0.65
0.73
- 62
1.97
1.85
1.00
28.00
20.00
- 35.00
0.53
0.81
- 52
1.97
1.85
1.00
28.00
20.00
- 35.00
0.53
0.81
- 52
1.81 2.62
1.61 2.29
1.00 1.00
27.00 44.00
13.00 26.00
- 30.00 - 45.00
0.45 0.74
0.61 0.72
- 51 - 67
2.40
2.16
1.00
39.00
25.00
- 42.00
0.67
0.76
- 62
2.18
2,02
1.00
33.00
23.00
- 39.00
0.61
0.80
- 58
2.02
1.89
1.00
29.00
21.00
- 36.00
0.55
0.82
- 53
1.88
1.88
1.00
20.00
20.00
- 36.00
0.51
1.00
- 47
1.95
1,88
1.00
27.00
22.00
- 35.00
0.53
0.88
- 51
1.67
1,61
1.00
20.00
16.00
- 28.00
0.40
0.87
- 42
1.86
1,86
1.00
23.00
23.00
- 34.00
0.51
1.00
- 46
1.77
1,71
1.00
22.00
18.00
- 31.00
0.45
0.88
- 45
1.70
1,52
1.00
24.00
11.00
-27.00
0.40
0.58
-47
1.69
1.58
1.00
22.00
14.00
- 28.00
0.41
0.74
- 45
2.00
2,00
1.00
26.00
26.00
- 37.00
0.57
1.00
- 50
2.06 2.02
1.84 1,80
1.00 1.00
32.00 31.00
18.00 17.00
- 36.00 - 35.00
0.55 0.53
0.69 0.68
- 57 - 56
1.81
1.61
1.00
27.00
13.00
-30.00
0.45
0.61
-51
2.26
1.60
1.00
47.00
4.00
- 35.00
0.58
0.15
- 69
2.03
1.78
1.00
32.00
16.00
- 35.00
0.53
0.64
- 57
2.51
2.36
1.00
38.00
30.00
- 45.00
0.73
0.87
- 62
2.82
2.65
1.00
44.00
35.00
- 49.00
0.82
0.88
- 67
1.84
1.81
1.00
23.00
21.00
- 33.00
0.49
0.95
- 46
1.84
1.67
1.00
27.00
15.00
- 31.00
0.46
0.67
- 51
1.84
1.74
1.00
25.00
18.00
- 32.00
0.48
0.81
- 49
2.24
2.17
1.00
32.00
28.00
- 41.00
0.65
0.92
- 57
1.97
1.94
1.00
26.00
24.00
- 36.00
0.55
0.95
- 50
2.26
1.85
1.00
40.00
15.00
- 38.00
0.60
0.52
- 64
1.77
1.64
1.00
24.00
15.00
- 30.00
0.44
0.74
- 48
1.92 2.19
1.79 1.92
1.00 1.00
27.00 36.00
18.00 19.00
- 34.00 -38.00
0.51 0.59
0.78 0.66
- 52 -60
1.79 2.25
1.56 1.97
1.00 1.00
27.00 37.00
11.00 20.00
- 29.00 - 39.00
0.43 0.61
0.54 0.67
- 51 - 61
1.98
1.77
1.00
31.00
17.00
- 34.00
0.52
0.67
- 55
S.R. Paterson et al. / Tectonophysics 247 (1995) 105-119
110
Table
1 (continued) Apparent
ellipsoids
V o l u m e loss
SI
LP
for X = 1
0.61
0.44
- 66
0.48
0.81
- 49
- 36.00
0.55
0.63
- 58
21.00
- 40.00
0.63
0.70
- 61
19.00
12.00
- 25.00
0.35
0.74
- 41
1.00
39.00
11.00
-35.00
0.55
0.41
63
2.05
1.00
32.00
25.00
- 39.00
0.61
0.86
- 56
2.32 2,27
2.10 2.07
1.00
37.00
24.00
41.00
0,65
0.76
- 61
1.00
36.00
24.00
40.00
0,63
0.77
- 60
1.86 1.83 1.84 1.84 1.64
1.86 1.67 1.65 1.65 1.64
1.00
23.00
23.00
- 34.00
0,51
1.00
- 46 - 50
1.00
18.00
2.36
2.17
1.00
37.00
2.15
1.97
1.00
33.00
22.00
2.18
2.12
1.00
31.00
27.00
1.94
1.77
1.00
28.00
2.03 2.15
1.86 1.95
1.00
2.10
1.92 1.87 1.94
Axial ratios
Apparent
X
Y
2.31
1.84
extensions (%)
Z
X
Y
Z
1.82
1.00
43.00
13.00
- 38.00
1.74
1.00
25.(10
18.00
32.00
2.08
1,81
1.00
33.00
16.00
2,28
2.02
1.00
37.00
1,59
1.49
1.00
2,14
1.71
2,16
New York Plateau
1.00
26.00
15.00
- 31.00
0,46
(/.70
1.00
27.00
14.00
- 31.00
0,46
0.65
51
1.00
27.00
14.00
-31.00
0,46
0.65
51
18.00
-28.00
0.40
1.00
39
26.00
- 42.00
0.67
0.81
- 61
- 38.00
0,59
0.77
- 57
- 40.00
0.63
0.92
- 56
17.00
34.00
0.51
0.73
- 53
30.00
19.00
- 36.00
0.55
0.75
55
1,00
33.00
21.00
- 38.00
0.59
0.75
- 58
1.92
1.00
32.00
21.00
- 37.00
0.57
0.76
- 56
1.79 1.71 1.70
1.00
27.00
18.00
- 34,00
0,51
0.78
- 52
1,00
27.00
16.00
32.00
0.48
0.71
-51
1.00
30.00
14.00
33.00
0.50
0.60
-55
2.13
2.05
1.00
30.00
25.00
- 39.00
0.60
0.90
- 55
2.02
1.91
1.00
29.00
22.00
- 36.00
0.55
0.84
- 53
1.95 1.83 1.94
1.77 1.60 1.83
1.00
29.00
17.00
- 34.00
0.51
0.71
- 54
1.00
28.00
12.00
- 30.00
0.45
0.56
52
1.00
28.00
21.00
- 34.00
0.52
0.83
51
1.85
1.71
1.00
26.00
16.00
32.00
0.47
0.73
- 50
2.23
1.95
2.85 2.20
2.50 2.02
1.00 1.00
36.00 48.00
19.0// 30.00
- 39.00 - 48.00
0.61 0.81
{/.67 0.75
61 - 69
1.00
34.00
23.00
- 39.00
0.61
0.78
- 58
2.85
2.48
1,00
48.00
29.00
- 48.00
0,80
0.74
- 69
2.55
2.23
1.00
43.00
25.00
- 44.00
0.72
0.71
- 66
2.36
2.05
1.00
39.00
21.00
- 41.00
0.65
0.68
- 63
2.85
2.48
1.00
48.00
29.00
- 48.00
0.80
0.74
- 69
2.20
2.02
1.00
34,00
23.00
- 39.00
0,61
0.78
- 58
2.85
2.48
.00
48.00
29.00
- 48.00
0.80
0.74
- 69
2.55
2.23
~.00
43.00
25.00
- 44.00
0.72
0.71
- 66
2.36 2.85
2.05 2.48
•0 0
39.00
21.00
- 41.00
0.65
0.68
- 63
•00
48.00
29.00
- 48.00
0.80
0.74
- 69
2.49 2.08
2.18 1.90
2.25 2.05
42.00 31,00 37.00
24.00 20.00 20.00
- 43.00 - 37.00 - 39.(10
0.70 0.57 0.61
0.70 0.76 0.67
- 65 - 56 - 61
1.93
1.97 1.88 1.72
~.00 •0 0 .(10 .00 •0 0
31.00 29.00
20.00 15.00
- 36.00 - 33.00
0.55 0.49
0,76 0,65
- 55 - 54
2.41 2.09
2.12 1.83
1.00 1.00
40.00 34.00
23.00 17.00
42.00 - 36.00
0.67 0.56
0,71 0,63
- 64 - 58
1.00 1.00
24.00 42.00
17.00 24.00
31.00 - 43,00
0.46 0.70
0,80 0.70
- 47 - 65
1.80
1.70
2.49
2.18
2.36
2.07
1.00
39.00
22.00
- 41.00
0.65
0.70
- 63
2.25
1.97
2.36
2.05
1.00 1.00
37.00 39.00
20.{)0 21.00
39.00 - 41.00
0.61 0.65
0.67 0.68
- 61 - 63
S.R. Paterson et al. / Tectonophysics 247 (1995) 105-119
111
Table 1 (continued) Apparent extensions (%)
Apparent ellipsoids Volume loss
X
Y
Z
X
Y
SI
LP
for X = 1
2.75 2.67 2.92 2.84 3.13 3.11 3.11 3.11 2.94 3.00 3.00 2.76 2.92 3.11 2.45 2.60
2.75 2.67 2.92 2.84 3.13 3.11 3.11 3.11 2.94 3.00 3.00 2.76 2.92 3.11 2.45 2.60
1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00
40.00 39.00 43.00 42.00 47.00 46.00 46.00 46.00 44.00 44.00 44.00 41.00 43.00 46.00 35.00 38.00
40.00 39.00 43.00 42.00 47.00 46.00 46.00 46.00 44.00 44.00 44.00 41.00 43.00 46.00 35.00 38.00
49.00 48.00 51.00 50.00 53.00 53.00 53.00 53.00 51.00 52.00 52.00 49.00 51.00 53.00 45.00 47.00
0.82 0.80 0.87 0.85 0.93 0.93 0.93 0.93 0.88 0.90 0.90 0.83 0.87 0.93 0.73 0.78
1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00
-
1.27 1.22 1.19 1.13 1.20 1.23 1.30 1.27
1.27 1.22 1.19 1.13 1.20 1.23 1.30 1.27
1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00
8.00 7.00 6.00 4.00 6.00 7.00 9.00 8.00
8.00 7.00 6.00 4.00 6.00 7.00 9.00 8.00
- 15.00 - 12.00 - 11.00 - 8.00 - 12.00 - 13.00 - 16.00 - 15.00
0.20 0.16 0.14 0.10 0.15 0.17 0.21 0.20
1.00 1.00
- 21 - 18 - 16 - 12 - 17 - 19 - 23 - 21
Axial ratios
Z
Yorkshire
72 PS shale Yorkshire-S Yorkshire-S
-
-
-
64 63 66 65 68 68 68 68 66 67 67 64 66 68 59 62
Green River Basin
rocks,
a f f e c t e d by t h e faulting.
visual comparison, each ellipsoid can be repres e n t e d by a single d o t o n a m o d i f i e d F l i n n diag r a m (e.g., F i g . 1) w h e r e t h e v e r t i c a l a n d h o r i z o n -
Fabric
The
ellipsoids
preferred
orientation
of
phyllosilicate
g r a i n s i n all o f t h e a b o v e s a m p l e s w a s d e t e r m i n e d u s i n g a n X - r a y g o n i o m e t e r in t r a n s m i s s i o n m o d e . Techniques of sample preparation, data collect i o n , a n d a n a l y s i s a r e s u m m a r i z e d in O e r t e l (1983). T h e s e p r e f e r r e d o r i e n t a t i o n s c a n b e u s e d t o d e f i n e a n e l l i p s o i d (e.g., W h e e l e r , 1986), w h i c h in t h i s p a p e r w e call t h e f a b r i c e l l i p s o i d . S i n c e o u r m a i n g o a l in t h i s p a p e r is t o c o m p a r e t h e s e e l l i p s o i d s t o e l l i p s o i d s m e a s u r e d in d e f o r m e d
graphically
1.00 1.00 1.00 1.00 1.00
Hildebrand-Mittlefehldt's research was investigation of the strain field caused by a nearby fault, only those measurements were used that are un-
3.
we
1.00
present
and
manipulate
t h e s e f a b r i c e l l i p s o i d s in a f a s h i o n s i m i l a r t o t h a t u s u a l l y d o n e w i t h s t r a i n e l l i p s o i d s ( T a b l e 1). F o r
tal a x e s a r e t h e n a t u r a l l o g a r i t h m s o f t h e X / Y a n d Y / Z r a t i o , r e s p e c t i v e l y , a n d w h e r e X , Y, a n d Z are the greatest, intermediate, and smallest principal axes of the fabric ellipsoid. The diagram a x e s m a y a l s o b e t a k e n t o r e p r e s e n t E1 - E 2 a n d E 2 - E 3 w h e r e E l , E2 , a n d E 3 a r e p r i n c i p a l n a t u r a l s t r a i n s ( R a m s a y a n d H u b e r , 1991). A g r e a t m a j o r i t y o f t h e f a b r i c e l l i p s o i d s fall a l o n g o r n e a r t h e h o r i z o n t a l axis o f t h e F l i n n d i a g r a m , t h a t is t h e l o c a t i o n o n t h i s d i a g r a m where ellipsoids are oblate and uniaxial. Small d e p a r t u r e s f r o m s t r i c t u n i a x i a l i t y , w h i c h will p l a c e
112
S.R. Paterson et al. / Tectonophysics 247 (1995) 105-119
the ellipsoid above the horizontal axis, could be due to either m e a s u r e m e n t errors or to real, local heterogeneities in the strain field caused by variations in compaction. In either case, if the ellipsoids were averaged (Oertel, 1981) they would result in a uniaxial oblate ellipsoid. Unfortunately, we cannot further evaluate these possibilities since many of our samples have poorly con-
strained X and Y axis orientations (e.g., an unknown amount of vertical axis rotation in core samples). Considerable independent evidence that pelitic rocks undergo compaction during burial (see reviews listed above), leads us to interpret the ellipsoids near the horizontal axis in Fig. 2 as reflecting such compaction. In that case, the sampled
Primary fabric ellipsoids from mudstones
2.6 2.4 2.2 2 1.8 1.6 ¢~ 1 . 4 uJ |
"~ 1.2 1 0.8 0.6 0.4
a a
0.2 0 0.2
0.4
0.6
0.8
1
1.2
1.4
1.6
1.8
2
2.2
2.4
2.6
E2-E3
•
DSDP 174A
[] Great Valley
* Pigeon Point
00DP
•
Labrador Trough
~ New York Plateau
•
o Green River Basin
Yorkshire
Hole 808c
Fig. 1. Modified (i.e., logarithmic) Flinn plot displaying clay fabric ellipsoids measured in compacted but tectonically weakly deformed to undeformed pelitic rocks. Sample locations and rock types decribed in text.
S.R. Paterson et al. / Tectonophysics 247 (1995) 105-119
113
mate of compaction with the possible exception of a few samples from the Labrador Trough. Somewhat surprising results come from the deep sea core samples located near or in accretionary wedges (Fig. 2). These samples generally have fabric ellipsoids with small ratios compared to most of those on Fig. 1 (note the change in axis scale from Fig. 1), and only a few plot near the horizontal axis (Fig. 2). Both Morgan and Karig (1993) and Behrmann and Kopf (1993) note that in this setting there is no simple relationship
units underwent compactions (i.e., water and volume loss causing vertical shortening) ranging from 3% to 74% (Table 1). Ellipsoids that plot far from this axis must include effects of other processes, the most likely of which is minor bedding-parallel shortening in the New York plateau and Labrador Trough samples (Oertel et al., 1989; Evans et al., 1989; Krishnan and Oertel, 1980). Because this shortening is small compared to the compaction in these samples, the bed thickening does not significantly affect our esti-
Primary fabric ellipsoids from deep sea mudstonss
0.5
0.4
0.3 I,IJ |
0.2 o
•
/
/
[]
0.1 o
[]
D
[]
0
I
I
0.1
0.2
I
I
0.3
0.4
0.5
E2-E3 •
DSDP 174a
[] ODP 808c
Fig. 2. Modified Flinn plot of clay fabrics in DSDP core 174A and O D P core 808C. ODP data from Morgan and Karig (1993). Note that these data come from oceanic sediments near or in accretionary wedges.
114
S.R. Paterson et al. / Tectonophysics 247 (1995) 105-119
bedding. Many of these ellipsoids plot on or near the horizontal axis on a Flinn diagram and reflect variable compactions (volume loss by vertical shortening) of 3% to 74% (Table 1). Both Curtis et al. (1980) and Baker et al. (1993) argued that the amount of compaction is to a large degree controlled by burial depth at shallow crustal levels (where initial porosities are high) and by the percent of nonplaty minerals (usually quartz and feldspar), a conclusion in good agreement with our data. In this regard, it is worthwhile to corn-
between depth of burial and compaction strain, a conclusion in agreement with our few data from Site 174A.
4. Discussion and conclusions
In most of the tectonically weakly to undeformed pelitic rocks measured in this study, the phyllosilicate grains define approximately uniaxial oblate ellipsoids with their X - Y planes parallel to
06
Comparison of adjacent sandstone and shale ellipsoids
0.5
0.4 -to w ,
A
0.3
A A 0.2
A Q
A
0.1
o
AA• 0
0.1
I 0.2
o
•
•
Q
•
I
I
0.3
0.4
•
I 0.5
0.6
E2-E3
• Deep S e a C o r e Sandstones
• Pigeon Point Sandstones
© Deep SeaCore • Pigeon Point Shales Shales
A Great Valley Sandstones
• Great Valley Shales
Fig. 3. Modified Flinn plot comparing fabric ellipsoids from adjacent sandy and shaly layers in the DSDP core 174A, Great Valley sequence, and the Pigeon Point Formation. Note that shales generally show a much greater degree of flattening. See text and Paterson and Yu (1994) for sample description.
S.R. Paterson et al. / Tectonophysics 247 (1995) 105-119
pare the phyllosilicate ellipsoids measured in this study to fabric ellipsoids measured by Paterson and Yu (1994) in sandy interbeds using the shape and orientation of quartz grains (Fig. 3). These data have been collected from immediately adjacent quartz-rich and phyllosilicate-rich beds in the Great Valley, Pigeon Point, and DSDP 174A areas. Quartz fabric ellipsoids have highly variable shapes and orientations and low axial ratios compared to the phyllosilicate ellipsoids (Fig. 3). The phyllosilicate ellipsoids tend to have larger ratios and nearly uniaxial oblate shapes, again indicating more compaction in the phyllosilicaterich layers. Thus after compaction, and presumably prior to tectonism, the pelitic rocks measured in this study start with a bedding-parallel foliation associated with values of shortening between 3% and 74%, whereas nearby sandy beds undergo relatively little compaction and have no foliation. Although well documented in the sedimentological literature, it remains surprising how often this result is downplayed in structural studies, particularly studies looking at the following: the timing and number of tectonic foliations in deformed rocks, the interpretation of early bedding-parallel foliations (Maltman, 1981 is a notable exception), the development of cleavage refraction, and the calculation of later tectonic strains. We return to the issue of estimating tectonic strains below. Limited results from DSDP and ODP cores suggest that clay fabric ellipsoids in samples from near accretionary wedges fall between uniaxial flattening and plane strain (Fig. 2). One obvious explanation for the triaxial shapes is that these samples have undergone both compaction and horizontal shortening. Morgan and Karig (1993) argued that their samples underwent on average 10% horizontal shortening perpendicular to the trench axis. They note that this shortening is particulary pronounced immediately above the decollement. Byrne et al. (1991) and O'Brien et al. (1993) have measured magnetic anisotropy ellipsoids from several DSDP sites and noted that they are triaxial and have their long axes perpendicular to plate motions. Lack of X and Y orientation determinations for our deep sea samples hampers testing this hypothesis, but the shapes of
115
our ellipsoids from site 174A are compatible with the suggestions of Byrne et al. (1991) and Morgan and Karig (1993). Paterson and Tobisch (1993) noted a similar relationship between fabric ellipsoid X axes and the inferred trench axis in shales from accreted turbidites along the California coast (Pigeon Point samples) but could not separate this from the possible effects of later folding. We have also found a similar relationship in the Great Valley samples in the present study, that is calculated X axes of the clay fabric ellipsoids are roughly parallel to the fore arc basin axis. But again this direction also is parallel to local fold axes and flow directions in these sediments (Paterson and Yu, 1994). The variable compaction (Fig. 1; Table 1), and in some sediments the variable shapes of fabric ellipsoids (Fig. 2), make estimates of later tectonic strain difficult. To illustrate these difficulties, we compare the above data to results from deformed pelitic rocks in two regions in the western metamorphic belt, Sierra Nevada, California. The first group (Fig. 4a) are from clay-rich layers in argillite-chert sequences of the Jurassic-Triassic Calaveras Complex, which is thought to consist of oceanic sediments underplated onto an accretionary wedge (e.g., Paterson and Sample, 1988). The second group (Fig. 4b) are from clayrich layers in turbidite sequences of the Late Jurassic Mariposa Formation deposited in intraarc basins (e.g., Bogen, 1984). Both these data sets come from strongly tectonically deformed rocks with well-developed, steeply dipping cleavages parallel to axial planes of folds. Bedding is usually subparallel to cleavage except in fold hinges. We argue that given the spread in ellipsoid shapes and intensities (i.e., distance from origin on Flinn diagram) of both tectonically deformed and undeformed data, it is difficult and commonly impossible to determine uniquely the amount of compaction, horizontal shortening, and later tectonic strain undergone by each sample. Given the data in Fig. 4, is it possible to put any constraints on likely strain paths? Since the Calaveras samples are accreted oceanic sediments, they probably underwent an early history of compaction and some horizontal
116
S.R. Paterson et el./Tectonophysics 247 (1995) 105-119 Strain
ellipsoids
from
Calaveras
Complex
2.6 2.4
2.2 2 1.8 1.6
m.
1.4
1
O.8 0.6
=l=q=•
0.4 O,2
•
II
0 0.00
0 . 2 0 0.40
0.60 0.80
1.00 1.20 1 . 4 0 1.60 1.80 2 . 0 0 2 . 2 0 2.40
2.60
E2-E3
Strain ellipsoids
from
Foothills
Terrane
2.6 2.4 2.2 2 1.8 1.6
=,
¢~ 1.4
~ 1.2 1 0.8
/
0.6
•" -
0.4 /
0.2 0
• .•
"1 0.2
=....::..
.
•
o:.,
•
•
w•
•
--
• .
"°tq.
I
?
+
I
I
I
[
I
I
I
I
0.4
0,6
0.8
1
1.2
1,4
1,6
1.8
2
2.2
2.4
E2-E3
2.6
S.R. Paterson et al. / Tectonophysics 247 (1995) 105-119
117
Summary of fabric ellipsoids data 2.6 2.4 2.2 2
1.8 1.6 lls UJ I
Intra-arc
Slates
1.4 1.2
1
Nondeforme/
0.8 0.6
Calaveras
Sandstones
Underplate
0.4
Slates 0.2
Compacted Shales 0 0
0.2
0.4
0.6
0.8
1
1.2
1.4
1.6
1.8
2
2.2
2.4
2.6
E2-E3 Fig. 5. Modified Flinn plot summarizing fields where fabric ellipsoids measured in different areas or rock types are located. Nondeformed sandstone ellipsoids are described by Paterson and Yu (1994).
shortening similar to the deep sea samples in Fig. 2. The simplest path continuation to end up with the data on Fig. 4a would require approximately equal amounts of plane strain and compaction, presumably occurring simultaneously during underplating. Bell (1985) and Baker et al. (1993) have described geometrically similar paths for deformed rocks from the Welsh slate belt.
The Mariposa samples (Fig. 4b) are more difficult to evaluate as a group because of the wide variability in final shapes. Ellipsoids near the plane strain line and in the constrictional field generally are located in ductile shear zones or fold hinges (Paterson et al., 1989). The remainder of the data come from fold limbs or regions of no folding. Ramsay and Wood (1973) and Bell (1985)
Fig. 4. Modified Flinn plots showing fabric ellipsoids from (a) clay-rich layers in argillite-chert sequences of the Jura-Triassic Calaveras Complex, Sierra Nevada, California; and (b) clay-rich layers in turbidite sequences of the Late Jurassic Mariposa Formation, Foothills Terrane, Sierra Nevada, California. The former are tectonically deformed samples of oceanic sediments underplated onto an accretionary wedge (Paterson and Sample, 1988), whereas the latter are tectonically deformed oceanic turbidites deposited in intra-arc basins (Bogen, 1984). For description of sample localities and measuring techniques see Paterson et al. (1989).
118
S.R. Paterson et al. / Tectonophysics 247 (1995) 105-119
have described paths in fold hinges or involving bedding parallel shortening by which compacted rocks can end up with plane strain or even with constrictional fabrics. Such paths may lead to the plane strain and constrictional data on Fig. 4b. Structural studies in the Foothills terrane generally support plane strain deformation (e.g., Paterson et al., 1989), and various amounts of plane strain plus compaction are certainly compatible with the remainder of the data in Fig. 4b. Yet, the most we can truly say is that the data are compatible with a variety of scenarios all of which involve some compaction and later tectonic strain. Figure 5 summarizes the different fields defined by fabric ellipsoids measured in undeformed sandstones, compacted shales, and deformed intra-arc and accretionary wedge slates. The data support the following conclusions: (1) A bedding-parallel foliation forms in pelitic rocks during compaction, which in our samples is associated with up to 74% shortening; (2) the amount of compaction is in part controlled by equidimensional grains (i.e., usually q u a r t z - f e l d s p a r content), thus interbedded sandy layers undergo little to no compaction (e.g., Paterson and Yu 1994); and (3) it is impossible to determine precisely the amounts of compaction, horizontal shortening, and later tectonic strain in deformed samples because of the variability of initial compactions. Our data, and results presented by others, do indicate that strain paths in pelitic rocks always include early compaction, which may be followed by an episode of simultaneous compaction and plane strain in accretionary wedges and by a variety of other types of strain in other settings. We see two immediate directions for future work attempting to characterize strain paths in pelitic rocks. If measurements of compaction strain can be correlated with quartz + feldspar contents, it may be possible to better estimate the amount of early compaction experienced by tectonically deformed samples. Attempts to do so will obviously be complicated by undercompaction caused by abnormal fluid pressures (e.g., Engelder and Oertel, 1985; Evans et al., 1989), compaction during or after tectonism (e.g., Bell, 1985; Chen and Oertel, 1989), and removal of
quartz and feldspar from deformed samples. The initial results from deep sea cores also indicate that additional measurements in a wide variety of deep sea settings, particularly near and in accretionary wedges, would be useful.
Acknowledgements Acknowledgments by Paterson are made to the Donors of The Petroleum Research Fund, administered by the American Chemical Society, and to the Faculty Initiative Research Fund grant, University of Southern California, for partial support of this research. Paterson thanks Dr. Oertel for his many years of help with measuring and interpreting fabric data, use of his X-ray goniometer, and wonderful dry wit and humor. Oertel acknowledges support from many National Science Foundation grants for partial support of this research. We thank Julia Morgan for permission to use the O D P data, Othmar Tobisch for discussions about strain and fabrics in rock, and Richard Bennett for a review of the manuscript.
References Baker, D.W., Chawla, K.S. and Krizek, R.J., 1993. Compacton fabrics of pelites: experimental consolidation of kaolinite and implications for analysis of strain in slate. J. Struct. Geol., 15: 1123-1137. Behrmann, J.H. and Kopf, A., 1993. Textures and microfabrics in fine-grained muds and mudstones from site 808, Nankai accretonary prism. In: I.A. Hill, A. Taira, J.V. Firth, et al. (Editors), Proceedings of the Ocean Drilling Program, 131: 141-158. Bell, A., 1985. Strain paths during slaty cleavage formation -the role of volume loss. J. Struct. Geol., 7: 563-568. Bennett, R.H., Bryant, W.R. and Hulbert, M.H., 1991. Microstructures of Fine-Grained Sediments: From Mud to Shale. Springer-Verlag, New York, NY, 582 pp. Bogen, N.L., 1984. Stratigraphic and sedimentologic evidence of a submarine island-arc volcano in the lower Mesozoic Penon Blanco and Jasper Point Formations, Mariposa County, California. Geol. Soc. Am. Bull., 96: 1322-1331. Byrne, T., Bruckmann, W. and Owens, W.H., 1991. Correlation of core-scale structural fabrics, acoustic properties and magnetic anisotropy in an accreting sedimentary sequence (ODP Leg 131, Nankai Trough). EOS, Trans. Am. Geophys. Union, 72: 535.
S.R. Paterson et al. / Tectonophysics 247 (1995) 105-119 Chen, R.T. and Oertel, G., 1989. Strain history of the Los Prietos syncline, Santa Maria basin, California: a case of post-tectonic compaction. J. Struct. Geol., 11,359-11,551. Curtis, C.D., Lipshie, S.R., Oertel, G. and Pearson, M.J., 1980. Clay orientation in some Upper Carboniferous mudrocks, its relationship to quartz content and some inferences about fissility, porosity, and compactional history. Sedimentology, 27: 333-339. Engelder, T. and Oertel, G., 1985. Correlation between abnormal pore pressure and tectonic jointing in the Devonian Catskill delta. Geology, 13: 863-866. Evans, K., Oertel, G. and Engelder, T., 1989 Appalachian stress study 2: Analysis of Devonian shale core: some implications for the nature of contemporary stress variations and Alleghanian deformation in Devonian rocks. J. Geophys. Res., 94: 7155-7170. Glen, R.A., 1990. Formation of thrusting in some Great Valley rocks near the Franciscan Complex, California, and its implications for the tectonic wedging hypothesis. Tectonics, 9: 1451-1477. Hildebrand-Mittlefehldt, N., 1980. Deformation near a fault termination, part II: a normal fault in shales. Tectonophysics, 64: 211-234. Howell, D.G. and Joyce, J.E., 1981. Field guide to the Upper Cretaceous Pigeon Point Formation. In: V. Frizzell (Editor), Cretaceous and Paleocene Turbidites, Central California Coast. Soc. Econ. Paleontol. Mineral., Pacific Coast Section, Field Trip, 6: 61-70. Ingersoll, R.V., 1979. Evolution of the Late Cretaceous forearc basin, northern and central California. Geol. Soc. Am. Bull., 90 (part 1): 813-826. Joyce, J.E., 1981. A deformational history of the Pigeon Point Formation. In: V. Frizell (Editor), Upper Cretaceous and Paleocene Turbidites, Central California Coast. Soc. Econ. Paleontol. Mineral., Pacific Coast Section, Field Trip, 6: 57-61. Krishnan, T.K. and Oertel, G., 1980. Aspects of strain history in folded sediments from the Schefferville Mining District, Labrador Trough, Canada. Tectonophysics, 64: 33-46. Kulm, L.D., von Huene, R., Duncan, J.R., Ingle, J.C., Kling Jr., S.A., Musich, L.F., Piper, D.J.W., Pratt, R.M., Schrader, H.J., Weser, O.F. and Wise Jr., S.W., 1973. Initial reports of the Deep Sea Drilling Project, 18. U.S. Government Printing Office, Washington D.C., pp. 97-168. Maltman, A.J., 1981. Primary bedding-parallel fabrics in structural geology. J. Geol. Soc. London, 134: 57-70. Margara, K., 1978. Compaction and Fluid Migration. Elsevier, Amsterdam, 320 pp. Morgan, J.K. and Karig, D.E., 1993. Ductile strains in clay-rich sediments from Hole 808C: preliminary results using X-ray pole figure goniometry. In: I.A. Hill, A. Taira, J.V. Firth et al. (Editors), Proceedings of the Ocean Drilling Program. U.S. Government Printing Office, Washington D.C., 131: 141-158. O'Brien, D.K., Manghnani, M.H., Tribble, J.S. and Wenk,
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H.R., 1993. Preferred orientation and velocity anisotropy in marine clay-bearing calcareous sediments. In: Microstructures in Limestones.,? Oertel, G., 1981. Strain estimation from scattered observations in an inhomogeneously deformed domain of rocks. Tectonophysics, 77: 133-150. Oertel, G., 1983. The relationship of strain and preferred orientation of phyllosilicate grains in rocks - - a review. Tectonophysics, 100: 413-447. Oertel, G. and Curtis, C.D., 1972. Clay-ironstone concretion preserving fabrics due to progressive compaction. Geol. Soc. Am. Bull., 83: 2597-2606. Oertel, G., Engelder, T. and Evans, K., 1989. A comparison of the strain of crinoid columnals with that of their enclosing silty and shaly matrix on the Appalachian Plateau, New York. J. Struct. Geol., 11: 975-993. Paterson, S.R. and Sample, J.C., 1988. The development of folds and cleavages in slate belts by underplating in accretionary complexes: a comparison of the Kodak Formation, Alaska and the Calaveras Complex, California. Tectonics, 7: 859-874. Paterson, S.R. and Tobisch, O.T., 1993. Pre-lithification structures, deformation mechanisms, and fabric ellipsoids in slumped turbidites from the Pigeon Point Formation, California. Tectonophysics, 222: 135-149. Paterson, S.R. and Yu, H., 1994. Primary fabric ellipsoids in sandstones: implications for depositional processes and strain analysis. J. Struct. Geol., 16: 505-518. Paterson, S.R., Tobisch, O.T. and Bhattacharyva, T., 1989. Regional structural and strain analyses of terranes in the Western Metamorphic Belt, Sierra Nevada, California. J. Struct. Geol., 11: 255-273. Ramsay, J.G. and Huber, M.I., 1987. Techniques of Modern Structural Geology; volume 2: Folds and Fractures. Academic Press, New York, NY, 307 pp. Ramsay, J.G. and Wood, D.S., 1973. The geometric effects of volume change during deformation processes. Tectonophysics, 16: 263-277. Reed, W.E. and Oertel, G., 1978. Is compaction a factor in organic diagenesis? Geol. Soc. Am. Bull., 89: 658-662. Rieke, H.H. and Chilingarian, G.V., 1974. Compaction of Argillaceous Sediments. Elsevier, Amsterdam, 424 pp. Tobisch, O.T., 1984. Development of foliation and fold interference patterns produced by sedimentary processes. Geology, 12: 441-444. Tyler, J.H., 1972. Pigeon Point Formation, an upper Cretaceous shoreline succession, central California coast. J. Sediment. Petrol., 42: 537-557. Wentworth, C.M., 1960. Sedimentary structures and inferred turbidity current origin of the Cretaceous Pigeon Point Formation, San Mateo County, California. M.S. thesis, Stanford Univ., Stanford, CA, 60 pp. Wheeler, J., 1986. Strain analysis in rocks with pretectonic fabrics. J. Struct. Geol., 8: 887-898.
Tectonophysics 247 (1995) 105-119
Primary and tectonic fabric intensities in mudrocks Scott R. Paterson a,., Hao Yu a, Gerhard Oertel b a Department of Earth Sciences, University of Southern California, Los Angeles, CA 90089-0740, USA b Department of Earth and Space Sciences, University of California, Los Angeles, Los Angeles, CA 90024-1567, USA Received 18 May 1994; revised version accepted 20 September 1994
Abstract
In this paper, we present a summary of quantitative measurements of the preferred orientation of phyllosilicate grains in a variety of tectonically weakly deformed to undeformed, fine-grained, pelitic rocks. We also briefly examine a small amount of data on clay fabrics in deep sea cores, and compare them to fabrics in deformed equivalents from the Mariposa Formation and Calaveras Complex, Sierra Nevada, California. The measured fabric ellipsoids support the following conclusions: (1) a bedding-parallel foliation forms in pelitic rocks during compaction that in our samples is associated with shortening between 3% and 74%; (2) the amount of compaction is in part controlled by the quartz-feldspar content: interbedded sandy layers undergo little to no compaction and do not develop bedding parallel foliations; and (3) it is impossible to determine precisely the amounts of compaction, horizontal shortening, and later tectonic strain in deformed samples because of the variability of initial compactions. Our data, and results presented by others, indicate that strain paths in pelitic rocks always include early compaction, followed by an episode of simultaneous compaction and plane strain in accretionary wedges, and by a variety of other types of strain in other settings.
1. Introduction
The shapes and preferred orientations of objects have been used extensively to evaluate strain in rocks. The preferred orientation of platy minerals has proven useful in this regard in finegrained sedimentary rocks, which often lack other suitable markers. However, there are a variety of processes that orient grains in sedimentary rocks including alignment during deposition, compaction, and pre- and post-lithifaction tectonic strain. Determining the magnitude of tectonic
* Corresponding author.
strain requires some information about prelithification fabrics. T h e r e exists a vast literature on the changes in microstructure and physical properties of muds and silts. For example, older summaries by Rieke and Chilingarian (1974) and Margara (1978) and recent work summarized in fifty-nine papers in Bennett et al. (1991) discuss the processes, microstructures, and resulting physical properties that occur during deposition, compaction, and lithification of fine-grained sediments. Baker et al. (1993) summarize changes in porosity, seismic velocity, seismic velocity anisotropy, and microstructure during compaction of pelites as well as new experimental results. Oertel (1983) has summarized techniques for measuring preferred orientations of phyllosili-
0040-1951/95/$09.50 © 1995 Elsevier Science B.V. All rights reserved SSDI 0 0 4 0 - 1 9 5 1 ( 9 4 ) 0 0 1 9 9 - 5
106
S.R. Paterson et al. / Tectonophysics 247 (1995) 105-119
cate grains and the validity of assumptions made in using the March model to relate clay fabrics to strain. These summaries and previous studies support the following: (1) compaction of pelitic rocks increases with depth during burial; (2) clay particles rotate towards parallelism with bedding during compaction; and (3) the magnitude of compaction is to a large degree controlled by the ratio of phyllosilicate minerals to more equidimensional grains (e.g., Curtis et al., 1980). Oertel (1983) and Baker et al. (1993) also concluded that the preferred orientation of phyllosilicates successfully reflects strain caused by compaction and tectonic processes. What is particularly useful to structural geologists, and much less commonly published, are quantitative measurements of the three-dimensional preferred orientation of phyllosilicate grains. Oertel and coworkers (i.e., Oertel and Curtis, 1972; Reed and Oertel, 1978; Curtis et al., 1980; Hildebrand-Mittlefehldt, 1980; Krishnan and Oertel, 1980; Engelder and Oertel, 1985; Evans et al., 1989; Oertel et al., 1989) have published some such data collected in rocks that at the scale of individual samples are not strongly deformed by tectonism. This paper presents a summary of data collected by Oertel and Paterson in a variety of tectonically weakly deformed to undeformed, fine-grained, pelitic rocks. We also examine a small amount of data on clay fabrics in deep sea cores, and compare them to fabrics in their deformed equivalents from the Mariposa Formation and Calaveras Complex, Sierra Nevada, California. Our results are compatible with the conclusions noted above. Our intention is not to reexamine these conclusions but, instead, provide a summary of quantitative measurements and to explore the implications of these measurements for studies of deformed equivalents.
2. Description of samples Deep Sea Core 174A: This core, drilled in the distal portion of the Astoria Fan off the coast of Oregon and seaward of the toe of the accre-
tionary prism, reached 879 m below the sea floor and recovered rocks from two units: (1) Upper Pleistocene thick- to thin-bedded, medium to fine turbiditic sands consisting of continental detritus deposited in the Astoria fan (0-284 m); and (2) the underlying Plio-Pleistocene thin-bedded sands, silts, and clays (284-879 m) representing abyssal plain deposits (Kulm et al., 1973). Ten samples were collected from the core, only four of which gave useful results (Table 1). During DSDP coring, an unknown amount of rotation around the vertical core axis occurred. Thus, orientations can only be measured with respect to horizontal. The preservation of horizontal bedding, laminations, and other sedimentary structures (Kulm et al., 1973) indicates that these samples were not significantly disrupted during coring or sample collection. Great Valley sequence: Six samples of claybearing siltstones were collected from a transect through the Cretaceous Great Valley sequence, California, in the Wilbur Springs Quadrangle. These interbedded sandstones and siltstones represent proximal and distal fore-arc fan sequences that after lithification have been tilted and locally deformed by bedding parallel faults (Ingersoll, 1979; Glen, 1990). However, at least in the sampled area, the rocks are not cleaved and show no microstructural evidence of crystal-plastic strain (Paterson and Yu, 1994). The sampled layers consist of 15-25% phyllosilicate grains, 50-75% quartz and feldspar grains, and small amounts of lithic fragments and volcanic detritus. Pigeon Point Formation: The Upper Cretaceous Pigeon Point Formation consists of turbiditic sequences of sandstone, siltstone, mudstone, and local conglomerate (Wentworth, 1960; Tyler, 1972). Facies F (see Howell and Joyce, 1981) is characterized by abundant slump structures. This facies consists of silt and sand distal turbidites with local, discontinuous mudstone and conglomerate layers. The mudstone consists of about 50% quartz, 10-20% feldspar, and variable amounts of micas, chlorite and clays. Organic material makes up a small but visible part in many samples, and some organic-rich layers are present. Although these rocks have been mildly deformed by open, upright, regional-scale folds
S.R. Paterson et aL / Tectonophysics 247 (1995) 105-119
and brittle faults (Howell and Joyce, 1981), metamorphic tectonites have not developed (Paterson and Tobisch, 1993), and the fabrics formed by penecontemporaneous slumping or compaction (e.g., Tobisch, 1984; Paterson and Tobisch, 1993). Eight samples have been analyzed from outcrops of facies F along a beach immediately south of Pigeon Point (Table 1). O D P H o l e 808c: To add to our data set of deep sea sediments we have included measured clay fabric orientatons from 9 samples collected by Morgan and Karig (1993) from ODP Hole 808c, drilled in the toe of the Nankai accretionary prism. Samples from depths ranging from 318 to 1137 m below the sea floor consist of hemipelagic silty-claystone with phyllosilicates making up 30% to 40%. They include both accreted rocks and subducted abyssal plain deposits. Behrmann and Kopf (1993) also measured clay fabrics in 11 samples from this site with samples obtained from depths of 23 to 533 m below the sea floor. Unfortunately, these authors did not present all three ratios of the calculated ellipsoids, thus precluding direct comparison of their results with data presented in this paper. Labrador Trough: These samples were collected in shaly beds of the Aphebian (early Proterozoic) exposed in the portion of the Labrador Trough that forms the Schefferville mining district of northeastern Canada (Krishnan and Oertel, 1980). They come mostly from the lowest stratigraphic unit in the trough, the Attikamagen Formation, some stratigraphic distance and at least one, or perhaps two, unconformities below the economically important Iron Formation. A few come from the Menihek Formation, which immediately overlies the Iron Formation. These calcareous, silty, and locally carbonaceous shales are deep water deposits formed in grabens in Archean basement. Whereas the central and eastern portions of the trough are moderately to strongly folded, its western rim is undeformed or only gently folded. The listed samples come mostly from this western margin of the trough or from long, straight, tectonically weakly strained limbs of moderate folds nearer the center of the trough. We excluded only those samples in which the shortest axis of the fabric ellipsoid (principal axis
107
of greatest contraction) departs significantly from the bedding-normal. This selection does not discriminate against some samples in which tectonic increments, even fairly large ones, happened to have been coaxial with the compaction strain. New York Plateau mudstones: Devonian mudstones were collected mostly from exposed, horizontal or gently dipping beds (maximum of 3°) of the Catskill Delta, New York Plateau, although a few came from borehole cores in the same region (Engelder and Oertel, 1985; Oertel et al., 1989; Evans et al., 1989). This delta is a large clastic wedge with both marine and fluvial sediments deposited on continental basement. Most show a small amount of measurable tectonic strain, caused by bedding-parallel shortening approximately 50-100 Ma after deposition (Engelder and Oertel, 1985). But this tectonic strain is only a small fraction of the strain caused by compaction during burial. Yorkshire: This group are shales and mudstones of Pennsylvanian (Westphalian A) age collected in tectonically undeformed horizontal beds in the vicinity of Penistone (near Sheffield), Yorkshire, England (Oertel and Curtis, 1972, and unpublished). These samples come from units deposited in shallow epicontinental basins on a Paleozoic basement. Two of the samples are seat earths (marked in Table 1 by an appended -S), the clay-rich substrata of coal seams. Green River Basin: Strain in this group was determined at various points in a single, large (0.6 x 0.4 x 0.3 m) sample (Hildebrand-Mittlefehldt, 1980) of gently east-dipping beds of the Middle Fork Tongue member of the Eocene Green River Formation, collected in the Uinta Basin near Utah State Highway 50. These thin shale layers, interbedded with marly limestones, are part of a thick sequence of lake beds. The measurements were performed by HildebrandMittlefehldt, using the same X-ray technique as for all the other fabric measurements discussed in this paper. The results, however, are unavailable as tabulated numerical values, and were, therefore, estimated from their graphic representation on Fig. 4 of the quoted paper (the scale is more than adequate for the required two significant figures of strain). Although the main objective of
S.R. Paterson et al. / Tectonophysics 247 (1995) 105-119
108 Table 1 Primary fabric data from mudstones Axial ratios
Apparent
extensions (%)
Y
Z
Apparent
ellipsoids
Volume
Sl
LP
for X = 1
X
Y
Z
X
OP1
1.88
1.61
1.00
29.97
11.30
- 30.87
0.47
0.51
- 54
OP6
1.13
1.10
1.00
5.10
2.31
- 6.99
0.09
0.56
- 14
OP7
1.02
1.01
1.00
0.99
0.00
- 0.99
0.01
0.00
OP10
1.71
1.24
1.00
33.10
GVS25
1.57
1.46
1.00
19.07
GVS27
1.46
1.30
1.00
17.92
GVS30
1.61
1.36
1.00
GVS32
1.54
1.36
1.00
GVS34
1.58
1.57
GVS36
1.31
P22 P23
DSDP Cores 174A
-3.48
-0.20
- 3
-22.16
0.38
-58
10.73
-24.16
0.34
0.68
-41
5.00
- 19.23
0.27
0.39
- 39
23.99
4.73
- 22.99
0.34
0.29
- 48
20.37
6.30
- 21.84
0.31
0.42
- 43
1.00
16.72
15.98
- 26.13
0.37
0.97
- 37
1.25
1.00
11.14
6.05
- 15.16
0.20
0.65
- 27
1.74
1.53
1.00
25.55
10.39
- 27.85
0.41
0.54
- 49
1.45
1.26
1.00
18.61
3.07
- 18.20
0.27
0.24
- 40
P24 P25
1.58 1.66
1.39 1.47
1.00 1.00
21.55 23.30
6.94 9.19
- 23.07 - 25.72
0.33 0.37
0.44 0.52
- 44 - 47
P26
1.20
1.18
1.00
6.86
5.08
- 10.95
0.14
0.82
- 18
OP6
1.13
1.10
1.00
5.10
2.31
-6.99
0.09
0.56
- 14
OP7
1.02
1.01
1.00
0.99
0.00
- 0.99
0.01
0.00
- 3
P21
1.88
1.61
1.00
29.97
11.30
- 30.87
0.47
0.51
- 54
Great Valley
Pigeon Point
ODP Hole 808C 1
1.27
1.19
1.00
10.49
3.69
- 12.71
0.17
0.46
- 26
2
1.33
1.19
1.00
14.39
1.79
- 14.11
0.20
0.19
- 33
3
1.53
1.40
1.00
18.80
8.60
- 22.50
0.32
0.58
- 40
4
1.27
1.09
1.00
14.27
- 2.52
10.22
0.17
- 0.32
- 33
5
1.35
1.24
1.00
13.67
4.55
- 15.85
0.22
0.44
- 32
6 7
1.23 1.32
1.11 1.24
1.00 1.00
10.57 12.34
0.24 4.94
-9.78 - 15.17
0.14 0.21
0.04 0.51
-26 - 29
8 9
1.23 1.37
1.21 1.28
1.00 1.00
7.81 13.71
6.11 6.21
-12.59 - 17.19
0.17 0.24
0.85 0.57
-20 - 32
3.13
3.11
1.00
47.00
46.00
- 53.00
0.93
0.99
- 68
2.92
2.92
1.00
43.00
43.00
- 51.00
0.87
1.00
- 66
3.24
3.17
1.00
49.00
46.00
-54.00
0.95
0.97
-70
3.20
3.20
1.00
47.00
47.00
- 54.00
0.95
1.00
- 69
2.98
2.55
1.00
52.00
30.00
- 49.00
0.84
0.71
- 71
3.55
3.30
1.00
56.00
45.00
- 56.00
1.00
0.88
- 74
2.90
2.59
1.00
48.00
32.00
- 49.00
0.83
(I.79
- 69
2.33
2.00
1.00
40.00
20.00
- 40.00
0.64
0.64
- 63
2.27 2.87 2.41
2.15 2.50 2.12
1.00 1.00 1.00
34.00 49.00 40.00
27.00 30.00 23.00
- 41.00 - 48.00 - 42.00
0.65 (I.81 0.67
0.87 0.74 0.71
- 58 - 70 - 64
2.67 2.26
2.37 1.85
1.00 1.00
44.00 40.00
28.00 15.00
- 46.00 - 38.0(I
0.76 0.60
0.76 0.52
- 67 - 64
4.27 2.16 2.43
3.39 2.03 1.90
1.00 1.00 1.00
75.00 32.00 46.00
39.00 24.00 14.00
- 59.00 - 39.00 - 40.00
1.10 0.61 0.65
0.68 0.84 0.44
- 81 - 57 - 68
Labrador Trough
loss
S.R. Paterson et al. / Tectonophysics 247 (1995) 105-119
109
Table 1 (continued) Axial ratios
Apparent extensions (%)
A p p a r e n t ellipsoids
V o l u m e loss
X
Y
Z
X
Y
Z
Sl
LP
for X = 1
2.02
1.88
1.00
29.00
20.00
- 36.00
0.55
0.79
- 54
2.11
1.98
1.00
31.00
23.00
- 38.00
0.59
0.83
- 56
2.35
1.95
1.00
41.00
17.00
- 40.00
0.64
0.56
- 65
2.16
1.94
1.00
34.00
20.00
-38.00
0.59
0.71
-59
New York Plateau
1.85
1.71
1.00
26.00
16.00
- 32.00
0.47
0.73
- 50
1.65
1.49
1.00
22.00
10.00
- 26.00
0.37
0.59
- 45
2.00
1.91
1.00
28.00
22.00
- 36.00
0.55
0.86
- 52
1.87
1.79
1.00
25.00
20.00
- 33.00
0.49
0.87
- 49
2.31 1.64
2.12 1.50
1.00 1.00
36.00 21.00
25.00 11.00
-41.00 - 26.00
0.65 0.37
0.80 0.65
-60 - 44
2.83
2.52
1.00
47.00
31.00
- 48.00
0.81
0.78
- 68
1.71 2.06
1.58 1.84
1.00 1.00
23.00 32.00
14.00 18.00
- 28.00 - 36.00
0.41 0.55
0.72 0.69
- 46 - 57 - 48
1.87
1.81
1.00
25.00
21.00
- 33.00
0.50
0.90
2.06
1.94
1.00
30.00
22.00
- 37.00
0.57
0.82
- 55
1.93
1.73
1.00
29.00
16.00
- 33.00
0.50
0.68
- 53
2.28
2.03
1.00
37.00
22.00
- 40.00
0.63
0.72
- 61
2.34
2.08
1.00
38.00
23.00
- 41.00
0.65
0.73
- 62
1.97
1.85
1.00
28.00
20.00
- 35.00
0.53
0.81
- 52
1.97
1.85
1.00
28.00
20.00
- 35.00
0.53
0.81
- 52
1.81 2.62
1.61 2.29
1.00 1.00
27.00 44.00
13.00 26.00
- 30.00 - 45.00
0.45 0.74
0.61 0.72
- 51 - 67
2.40
2.16
1.00
39.00
25.00
- 42.00
0.67
0.76
- 62
2.18
2,02
1.00
33.00
23.00
- 39.00
0.61
0.80
- 58
2.02
1.89
1.00
29.00
21.00
- 36.00
0.55
0.82
- 53
1.88
1.88
1.00
20.00
20.00
- 36.00
0.51
1.00
- 47
1.95
1,88
1.00
27.00
22.00
- 35.00
0.53
0.88
- 51
1.67
1,61
1.00
20.00
16.00
- 28.00
0.40
0.87
- 42
1.86
1,86
1.00
23.00
23.00
- 34.00
0.51
1.00
- 46
1.77
1,71
1.00
22.00
18.00
- 31.00
0.45
0.88
- 45
1.70
1,52
1.00
24.00
11.00
-27.00
0.40
0.58
-47
1.69
1.58
1.00
22.00
14.00
- 28.00
0.41
0.74
- 45
2.00
2,00
1.00
26.00
26.00
- 37.00
0.57
1.00
- 50
2.06 2.02
1.84 1,80
1.00 1.00
32.00 31.00
18.00 17.00
- 36.00 - 35.00
0.55 0.53
0.69 0.68
- 57 - 56
1.81
1.61
1.00
27.00
13.00
-30.00
0.45
0.61
-51
2.26
1.60
1.00
47.00
4.00
- 35.00
0.58
0.15
- 69
2.03
1.78
1.00
32.00
16.00
- 35.00
0.53
0.64
- 57
2.51
2.36
1.00
38.00
30.00
- 45.00
0.73
0.87
- 62
2.82
2.65
1.00
44.00
35.00
- 49.00
0.82
0.88
- 67
1.84
1.81
1.00
23.00
21.00
- 33.00
0.49
0.95
- 46
1.84
1.67
1.00
27.00
15.00
- 31.00
0.46
0.67
- 51
1.84
1.74
1.00
25.00
18.00
- 32.00
0.48
0.81
- 49
2.24
2.17
1.00
32.00
28.00
- 41.00
0.65
0.92
- 57
1.97
1.94
1.00
26.00
24.00
- 36.00
0.55
0.95
- 50
2.26
1.85
1.00
40.00
15.00
- 38.00
0.60
0.52
- 64
1.77
1.64
1.00
24.00
15.00
- 30.00
0.44
0.74
- 48
1.92 2.19
1.79 1.92
1.00 1.00
27.00 36.00
18.00 19.00
- 34.00 -38.00
0.51 0.59
0.78 0.66
- 52 -60
1.79 2.25
1.56 1.97
1.00 1.00
27.00 37.00
11.00 20.00
- 29.00 - 39.00
0.43 0.61
0.54 0.67
- 51 - 61
1.98
1.77
1.00
31.00
17.00
- 34.00
0.52
0.67
- 55
S.R. Paterson et al. / Tectonophysics 247 (1995) 105-119
110
Table
1 (continued) Apparent
ellipsoids
V o l u m e loss
SI
LP
for X = 1
0.61
0.44
- 66
0.48
0.81
- 49
- 36.00
0.55
0.63
- 58
21.00
- 40.00
0.63
0.70
- 61
19.00
12.00
- 25.00
0.35
0.74
- 41
1.00
39.00
11.00
-35.00
0.55
0.41
63
2.05
1.00
32.00
25.00
- 39.00
0.61
0.86
- 56
2.32 2,27
2.10 2.07
1.00
37.00
24.00
41.00
0,65
0.76
- 61
1.00
36.00
24.00
40.00
0,63
0.77
- 60
1.86 1.83 1.84 1.84 1.64
1.86 1.67 1.65 1.65 1.64
1.00
23.00
23.00
- 34.00
0,51
1.00
- 46 - 50
1.00
18.00
2.36
2.17
1.00
37.00
2.15
1.97
1.00
33.00
22.00
2.18
2.12
1.00
31.00
27.00
1.94
1.77
1.00
28.00
2.03 2.15
1.86 1.95
1.00
2.10
1.92 1.87 1.94
Axial ratios
Apparent
X
Y
2.31
1.84
extensions (%)
Z
X
Y
Z
1.82
1.00
43.00
13.00
- 38.00
1.74
1.00
25.(10
18.00
32.00
2.08
1,81
1.00
33.00
16.00
2,28
2.02
1.00
37.00
1,59
1.49
1.00
2,14
1.71
2,16
New York Plateau
1.00
26.00
15.00
- 31.00
0,46
(/.70
1.00
27.00
14.00
- 31.00
0,46
0.65
51
1.00
27.00
14.00
-31.00
0,46
0.65
51
18.00
-28.00
0.40
1.00
39
26.00
- 42.00
0.67
0.81
- 61
- 38.00
0,59
0.77
- 57
- 40.00
0.63
0.92
- 56
17.00
34.00
0.51
0.73
- 53
30.00
19.00
- 36.00
0.55
0.75
55
1,00
33.00
21.00
- 38.00
0.59
0.75
- 58
1.92
1.00
32.00
21.00
- 37.00
0.57
0.76
- 56
1.79 1.71 1.70
1.00
27.00
18.00
- 34,00
0,51
0.78
- 52
1,00
27.00
16.00
32.00
0.48
0.71
-51
1.00
30.00
14.00
33.00
0.50
0.60
-55
2.13
2.05
1.00
30.00
25.00
- 39.00
0.60
0.90
- 55
2.02
1.91
1.00
29.00
22.00
- 36.00
0.55
0.84
- 53
1.95 1.83 1.94
1.77 1.60 1.83
1.00
29.00
17.00
- 34.00
0.51
0.71
- 54
1.00
28.00
12.00
- 30.00
0.45
0.56
52
1.00
28.00
21.00
- 34.00
0.52
0.83
51
1.85
1.71
1.00
26.00
16.00
32.00
0.47
0.73
- 50
2.23
1.95
2.85 2.20
2.50 2.02
1.00 1.00
36.00 48.00
19.0// 30.00
- 39.00 - 48.00
0.61 0.81
{/.67 0.75
61 - 69
1.00
34.00
23.00
- 39.00
0.61
0.78
- 58
2.85
2.48
1,00
48.00
29.00
- 48.00
0,80
0.74
- 69
2.55
2.23
1.00
43.00
25.00
- 44.00
0.72
0.71
- 66
2.36
2.05
1.00
39.00
21.00
- 41.00
0.65
0.68
- 63
2.85
2.48
1.00
48.00
29.00
- 48.00
0.80
0.74
- 69
2.20
2.02
1.00
34,00
23.00
- 39.00
0,61
0.78
- 58
2.85
2.48
.00
48.00
29.00
- 48.00
0.80
0.74
- 69
2.55
2.23
~.00
43.00
25.00
- 44.00
0.72
0.71
- 66
2.36 2.85
2.05 2.48
•0 0
39.00
21.00
- 41.00
0.65
0.68
- 63
•00
48.00
29.00
- 48.00
0.80
0.74
- 69
2.49 2.08
2.18 1.90
2.25 2.05
42.00 31,00 37.00
24.00 20.00 20.00
- 43.00 - 37.00 - 39.(10
0.70 0.57 0.61
0.70 0.76 0.67
- 65 - 56 - 61
1.93
1.97 1.88 1.72
~.00 •0 0 .(10 .00 •0 0
31.00 29.00
20.00 15.00
- 36.00 - 33.00
0.55 0.49
0,76 0,65
- 55 - 54
2.41 2.09
2.12 1.83
1.00 1.00
40.00 34.00
23.00 17.00
42.00 - 36.00
0.67 0.56
0,71 0,63
- 64 - 58
1.00 1.00
24.00 42.00
17.00 24.00
31.00 - 43,00
0.46 0.70
0,80 0.70
- 47 - 65
1.80
1.70
2.49
2.18
2.36
2.07
1.00
39.00
22.00
- 41.00
0.65
0.70
- 63
2.25
1.97
2.36
2.05
1.00 1.00
37.00 39.00
20.{)0 21.00
39.00 - 41.00
0.61 0.65
0.67 0.68
- 61 - 63
S.R. Paterson et al. / Tectonophysics 247 (1995) 105-119
111
Table 1 (continued) Apparent extensions (%)
Apparent ellipsoids Volume loss
X
Y
Z
X
Y
SI
LP
for X = 1
2.75 2.67 2.92 2.84 3.13 3.11 3.11 3.11 2.94 3.00 3.00 2.76 2.92 3.11 2.45 2.60
2.75 2.67 2.92 2.84 3.13 3.11 3.11 3.11 2.94 3.00 3.00 2.76 2.92 3.11 2.45 2.60
1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00
40.00 39.00 43.00 42.00 47.00 46.00 46.00 46.00 44.00 44.00 44.00 41.00 43.00 46.00 35.00 38.00
40.00 39.00 43.00 42.00 47.00 46.00 46.00 46.00 44.00 44.00 44.00 41.00 43.00 46.00 35.00 38.00
49.00 48.00 51.00 50.00 53.00 53.00 53.00 53.00 51.00 52.00 52.00 49.00 51.00 53.00 45.00 47.00
0.82 0.80 0.87 0.85 0.93 0.93 0.93 0.93 0.88 0.90 0.90 0.83 0.87 0.93 0.73 0.78
1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00
-
1.27 1.22 1.19 1.13 1.20 1.23 1.30 1.27
1.27 1.22 1.19 1.13 1.20 1.23 1.30 1.27
1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00
8.00 7.00 6.00 4.00 6.00 7.00 9.00 8.00
8.00 7.00 6.00 4.00 6.00 7.00 9.00 8.00
- 15.00 - 12.00 - 11.00 - 8.00 - 12.00 - 13.00 - 16.00 - 15.00
0.20 0.16 0.14 0.10 0.15 0.17 0.21 0.20
1.00 1.00
- 21 - 18 - 16 - 12 - 17 - 19 - 23 - 21
Axial ratios
Z
Yorkshire
72 PS shale Yorkshire-S Yorkshire-S
-
-
-
64 63 66 65 68 68 68 68 66 67 67 64 66 68 59 62
Green River Basin
rocks,
a f f e c t e d by t h e faulting.
visual comparison, each ellipsoid can be repres e n t e d by a single d o t o n a m o d i f i e d F l i n n diag r a m (e.g., F i g . 1) w h e r e t h e v e r t i c a l a n d h o r i z o n -
Fabric
The
ellipsoids
preferred
orientation
of
phyllosilicate
g r a i n s i n all o f t h e a b o v e s a m p l e s w a s d e t e r m i n e d u s i n g a n X - r a y g o n i o m e t e r in t r a n s m i s s i o n m o d e . Techniques of sample preparation, data collect i o n , a n d a n a l y s i s a r e s u m m a r i z e d in O e r t e l (1983). T h e s e p r e f e r r e d o r i e n t a t i o n s c a n b e u s e d t o d e f i n e a n e l l i p s o i d (e.g., W h e e l e r , 1986), w h i c h in t h i s p a p e r w e call t h e f a b r i c e l l i p s o i d . S i n c e o u r m a i n g o a l in t h i s p a p e r is t o c o m p a r e t h e s e e l l i p s o i d s t o e l l i p s o i d s m e a s u r e d in d e f o r m e d
graphically
1.00 1.00 1.00 1.00 1.00
Hildebrand-Mittlefehldt's research was investigation of the strain field caused by a nearby fault, only those measurements were used that are un-
3.
we
1.00
present
and
manipulate
t h e s e f a b r i c e l l i p s o i d s in a f a s h i o n s i m i l a r t o t h a t u s u a l l y d o n e w i t h s t r a i n e l l i p s o i d s ( T a b l e 1). F o r
tal a x e s a r e t h e n a t u r a l l o g a r i t h m s o f t h e X / Y a n d Y / Z r a t i o , r e s p e c t i v e l y , a n d w h e r e X , Y, a n d Z are the greatest, intermediate, and smallest principal axes of the fabric ellipsoid. The diagram a x e s m a y a l s o b e t a k e n t o r e p r e s e n t E1 - E 2 a n d E 2 - E 3 w h e r e E l , E2 , a n d E 3 a r e p r i n c i p a l n a t u r a l s t r a i n s ( R a m s a y a n d H u b e r , 1991). A g r e a t m a j o r i t y o f t h e f a b r i c e l l i p s o i d s fall a l o n g o r n e a r t h e h o r i z o n t a l axis o f t h e F l i n n d i a g r a m , t h a t is t h e l o c a t i o n o n t h i s d i a g r a m where ellipsoids are oblate and uniaxial. Small d e p a r t u r e s f r o m s t r i c t u n i a x i a l i t y , w h i c h will p l a c e
112
S.R. Paterson et al. / Tectonophysics 247 (1995) 105-119
the ellipsoid above the horizontal axis, could be due to either m e a s u r e m e n t errors or to real, local heterogeneities in the strain field caused by variations in compaction. In either case, if the ellipsoids were averaged (Oertel, 1981) they would result in a uniaxial oblate ellipsoid. Unfortunately, we cannot further evaluate these possibilities since many of our samples have poorly con-
strained X and Y axis orientations (e.g., an unknown amount of vertical axis rotation in core samples). Considerable independent evidence that pelitic rocks undergo compaction during burial (see reviews listed above), leads us to interpret the ellipsoids near the horizontal axis in Fig. 2 as reflecting such compaction. In that case, the sampled
Primary fabric ellipsoids from mudstones
2.6 2.4 2.2 2 1.8 1.6 ¢~ 1 . 4 uJ |
"~ 1.2 1 0.8 0.6 0.4
a a
0.2 0 0.2
0.4
0.6
0.8
1
1.2
1.4
1.6
1.8
2
2.2
2.4
2.6
E2-E3
•
DSDP 174A
[] Great Valley
* Pigeon Point
00DP
•
Labrador Trough
~ New York Plateau
•
o Green River Basin
Yorkshire
Hole 808c
Fig. 1. Modified (i.e., logarithmic) Flinn plot displaying clay fabric ellipsoids measured in compacted but tectonically weakly deformed to undeformed pelitic rocks. Sample locations and rock types decribed in text.
S.R. Paterson et al. / Tectonophysics 247 (1995) 105-119
113
mate of compaction with the possible exception of a few samples from the Labrador Trough. Somewhat surprising results come from the deep sea core samples located near or in accretionary wedges (Fig. 2). These samples generally have fabric ellipsoids with small ratios compared to most of those on Fig. 1 (note the change in axis scale from Fig. 1), and only a few plot near the horizontal axis (Fig. 2). Both Morgan and Karig (1993) and Behrmann and Kopf (1993) note that in this setting there is no simple relationship
units underwent compactions (i.e., water and volume loss causing vertical shortening) ranging from 3% to 74% (Table 1). Ellipsoids that plot far from this axis must include effects of other processes, the most likely of which is minor bedding-parallel shortening in the New York plateau and Labrador Trough samples (Oertel et al., 1989; Evans et al., 1989; Krishnan and Oertel, 1980). Because this shortening is small compared to the compaction in these samples, the bed thickening does not significantly affect our esti-
Primary fabric ellipsoids from deep sea mudstonss
0.5
0.4
0.3 I,IJ |
0.2 o
•
/
/
[]
0.1 o
[]
D
[]
0
I
I
0.1
0.2
I
I
0.3
0.4
0.5
E2-E3 •
DSDP 174a
[] ODP 808c
Fig. 2. Modified Flinn plot of clay fabrics in DSDP core 174A and O D P core 808C. ODP data from Morgan and Karig (1993). Note that these data come from oceanic sediments near or in accretionary wedges.
114
S.R. Paterson et al. / Tectonophysics 247 (1995) 105-119
bedding. Many of these ellipsoids plot on or near the horizontal axis on a Flinn diagram and reflect variable compactions (volume loss by vertical shortening) of 3% to 74% (Table 1). Both Curtis et al. (1980) and Baker et al. (1993) argued that the amount of compaction is to a large degree controlled by burial depth at shallow crustal levels (where initial porosities are high) and by the percent of nonplaty minerals (usually quartz and feldspar), a conclusion in good agreement with our data. In this regard, it is worthwhile to corn-
between depth of burial and compaction strain, a conclusion in agreement with our few data from Site 174A.
4. Discussion and conclusions
In most of the tectonically weakly to undeformed pelitic rocks measured in this study, the phyllosilicate grains define approximately uniaxial oblate ellipsoids with their X - Y planes parallel to
06
Comparison of adjacent sandstone and shale ellipsoids
0.5
0.4 -to w ,
A
0.3
A A 0.2
A Q
A
0.1
o
AA• 0
0.1
I 0.2
o
•
•
Q
•
I
I
0.3
0.4
•
I 0.5
0.6
E2-E3
• Deep S e a C o r e Sandstones
• Pigeon Point Sandstones
© Deep SeaCore • Pigeon Point Shales Shales
A Great Valley Sandstones
• Great Valley Shales
Fig. 3. Modified Flinn plot comparing fabric ellipsoids from adjacent sandy and shaly layers in the DSDP core 174A, Great Valley sequence, and the Pigeon Point Formation. Note that shales generally show a much greater degree of flattening. See text and Paterson and Yu (1994) for sample description.
S.R. Paterson et al. / Tectonophysics 247 (1995) 105-119
pare the phyllosilicate ellipsoids measured in this study to fabric ellipsoids measured by Paterson and Yu (1994) in sandy interbeds using the shape and orientation of quartz grains (Fig. 3). These data have been collected from immediately adjacent quartz-rich and phyllosilicate-rich beds in the Great Valley, Pigeon Point, and DSDP 174A areas. Quartz fabric ellipsoids have highly variable shapes and orientations and low axial ratios compared to the phyllosilicate ellipsoids (Fig. 3). The phyllosilicate ellipsoids tend to have larger ratios and nearly uniaxial oblate shapes, again indicating more compaction in the phyllosilicaterich layers. Thus after compaction, and presumably prior to tectonism, the pelitic rocks measured in this study start with a bedding-parallel foliation associated with values of shortening between 3% and 74%, whereas nearby sandy beds undergo relatively little compaction and have no foliation. Although well documented in the sedimentological literature, it remains surprising how often this result is downplayed in structural studies, particularly studies looking at the following: the timing and number of tectonic foliations in deformed rocks, the interpretation of early bedding-parallel foliations (Maltman, 1981 is a notable exception), the development of cleavage refraction, and the calculation of later tectonic strains. We return to the issue of estimating tectonic strains below. Limited results from DSDP and ODP cores suggest that clay fabric ellipsoids in samples from near accretionary wedges fall between uniaxial flattening and plane strain (Fig. 2). One obvious explanation for the triaxial shapes is that these samples have undergone both compaction and horizontal shortening. Morgan and Karig (1993) argued that their samples underwent on average 10% horizontal shortening perpendicular to the trench axis. They note that this shortening is particulary pronounced immediately above the decollement. Byrne et al. (1991) and O'Brien et al. (1993) have measured magnetic anisotropy ellipsoids from several DSDP sites and noted that they are triaxial and have their long axes perpendicular to plate motions. Lack of X and Y orientation determinations for our deep sea samples hampers testing this hypothesis, but the shapes of
115
our ellipsoids from site 174A are compatible with the suggestions of Byrne et al. (1991) and Morgan and Karig (1993). Paterson and Tobisch (1993) noted a similar relationship between fabric ellipsoid X axes and the inferred trench axis in shales from accreted turbidites along the California coast (Pigeon Point samples) but could not separate this from the possible effects of later folding. We have also found a similar relationship in the Great Valley samples in the present study, that is calculated X axes of the clay fabric ellipsoids are roughly parallel to the fore arc basin axis. But again this direction also is parallel to local fold axes and flow directions in these sediments (Paterson and Yu, 1994). The variable compaction (Fig. 1; Table 1), and in some sediments the variable shapes of fabric ellipsoids (Fig. 2), make estimates of later tectonic strain difficult. To illustrate these difficulties, we compare the above data to results from deformed pelitic rocks in two regions in the western metamorphic belt, Sierra Nevada, California. The first group (Fig. 4a) are from clay-rich layers in argillite-chert sequences of the Jurassic-Triassic Calaveras Complex, which is thought to consist of oceanic sediments underplated onto an accretionary wedge (e.g., Paterson and Sample, 1988). The second group (Fig. 4b) are from clayrich layers in turbidite sequences of the Late Jurassic Mariposa Formation deposited in intraarc basins (e.g., Bogen, 1984). Both these data sets come from strongly tectonically deformed rocks with well-developed, steeply dipping cleavages parallel to axial planes of folds. Bedding is usually subparallel to cleavage except in fold hinges. We argue that given the spread in ellipsoid shapes and intensities (i.e., distance from origin on Flinn diagram) of both tectonically deformed and undeformed data, it is difficult and commonly impossible to determine uniquely the amount of compaction, horizontal shortening, and later tectonic strain undergone by each sample. Given the data in Fig. 4, is it possible to put any constraints on likely strain paths? Since the Calaveras samples are accreted oceanic sediments, they probably underwent an early history of compaction and some horizontal
116
S.R. Paterson et el./Tectonophysics 247 (1995) 105-119 Strain
ellipsoids
from
Calaveras
Complex
2.6 2.4
2.2 2 1.8 1.6
m.
1.4
1
O.8 0.6
=l=q=•
0.4 O,2
•
II
0 0.00
0 . 2 0 0.40
0.60 0.80
1.00 1.20 1 . 4 0 1.60 1.80 2 . 0 0 2 . 2 0 2.40
2.60
E2-E3
Strain ellipsoids
from
Foothills
Terrane
2.6 2.4 2.2 2 1.8 1.6
=,
¢~ 1.4
~ 1.2 1 0.8
/
0.6
•" -
0.4 /
0.2 0
• .•
"1 0.2
=....::..
.
•
o:.,
•
•
w•
•
--
• .
"°tq.
I
?
+
I
I
I
[
I
I
I
I
0.4
0,6
0.8
1
1.2
1,4
1,6
1.8
2
2.2
2.4
E2-E3
2.6
S.R. Paterson et al. / Tectonophysics 247 (1995) 105-119
117
Summary of fabric ellipsoids data 2.6 2.4 2.2 2
1.8 1.6 lls UJ I
Intra-arc
Slates
1.4 1.2
1
Nondeforme/
0.8 0.6
Calaveras
Sandstones
Underplate
0.4
Slates 0.2
Compacted Shales 0 0
0.2
0.4
0.6
0.8
1
1.2
1.4
1.6
1.8
2
2.2
2.4
2.6
E2-E3 Fig. 5. Modified Flinn plot summarizing fields where fabric ellipsoids measured in different areas or rock types are located. Nondeformed sandstone ellipsoids are described by Paterson and Yu (1994).
shortening similar to the deep sea samples in Fig. 2. The simplest path continuation to end up with the data on Fig. 4a would require approximately equal amounts of plane strain and compaction, presumably occurring simultaneously during underplating. Bell (1985) and Baker et al. (1993) have described geometrically similar paths for deformed rocks from the Welsh slate belt.
The Mariposa samples (Fig. 4b) are more difficult to evaluate as a group because of the wide variability in final shapes. Ellipsoids near the plane strain line and in the constrictional field generally are located in ductile shear zones or fold hinges (Paterson et al., 1989). The remainder of the data come from fold limbs or regions of no folding. Ramsay and Wood (1973) and Bell (1985)
Fig. 4. Modified Flinn plots showing fabric ellipsoids from (a) clay-rich layers in argillite-chert sequences of the Jura-Triassic Calaveras Complex, Sierra Nevada, California; and (b) clay-rich layers in turbidite sequences of the Late Jurassic Mariposa Formation, Foothills Terrane, Sierra Nevada, California. The former are tectonically deformed samples of oceanic sediments underplated onto an accretionary wedge (Paterson and Sample, 1988), whereas the latter are tectonically deformed oceanic turbidites deposited in intra-arc basins (Bogen, 1984). For description of sample localities and measuring techniques see Paterson et al. (1989).
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have described paths in fold hinges or involving bedding parallel shortening by which compacted rocks can end up with plane strain or even with constrictional fabrics. Such paths may lead to the plane strain and constrictional data on Fig. 4b. Structural studies in the Foothills terrane generally support plane strain deformation (e.g., Paterson et al., 1989), and various amounts of plane strain plus compaction are certainly compatible with the remainder of the data in Fig. 4b. Yet, the most we can truly say is that the data are compatible with a variety of scenarios all of which involve some compaction and later tectonic strain. Figure 5 summarizes the different fields defined by fabric ellipsoids measured in undeformed sandstones, compacted shales, and deformed intra-arc and accretionary wedge slates. The data support the following conclusions: (1) A bedding-parallel foliation forms in pelitic rocks during compaction, which in our samples is associated with up to 74% shortening; (2) the amount of compaction is in part controlled by equidimensional grains (i.e., usually q u a r t z - f e l d s p a r content), thus interbedded sandy layers undergo little to no compaction (e.g., Paterson and Yu 1994); and (3) it is impossible to determine precisely the amounts of compaction, horizontal shortening, and later tectonic strain in deformed samples because of the variability of initial compactions. Our data, and results presented by others, do indicate that strain paths in pelitic rocks always include early compaction, which may be followed by an episode of simultaneous compaction and plane strain in accretionary wedges and by a variety of other types of strain in other settings. We see two immediate directions for future work attempting to characterize strain paths in pelitic rocks. If measurements of compaction strain can be correlated with quartz + feldspar contents, it may be possible to better estimate the amount of early compaction experienced by tectonically deformed samples. Attempts to do so will obviously be complicated by undercompaction caused by abnormal fluid pressures (e.g., Engelder and Oertel, 1985; Evans et al., 1989), compaction during or after tectonism (e.g., Bell, 1985; Chen and Oertel, 1989), and removal of
quartz and feldspar from deformed samples. The initial results from deep sea cores also indicate that additional measurements in a wide variety of deep sea settings, particularly near and in accretionary wedges, would be useful.
Acknowledgements Acknowledgments by Paterson are made to the Donors of The Petroleum Research Fund, administered by the American Chemical Society, and to the Faculty Initiative Research Fund grant, University of Southern California, for partial support of this research. Paterson thanks Dr. Oertel for his many years of help with measuring and interpreting fabric data, use of his X-ray goniometer, and wonderful dry wit and humor. Oertel acknowledges support from many National Science Foundation grants for partial support of this research. We thank Julia Morgan for permission to use the O D P data, Othmar Tobisch for discussions about strain and fabrics in rock, and Richard Bennett for a review of the manuscript.
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