- Email: [email protected]
IODP Sites 1256, U1309, and U1415
Chapter 4.2.1
Formation and Evolution of Oceanic Lithosphere: New Insights on Crustal Structure and Igneous Geochemistry from ODP/IODP Sites 1256, U1309, and U1415 Benoit Ildefonse,1,* Natsue Abe,2 Marguerite Godard,1 Antony Morris,3 Damon A.H. Teagle4 and Susumu Umino5 1Géosciences Montpellier, Université Montpellier, Montpellier, France; 2Institute for Research on Earth Evolution, Japan Agency for Marine-Earth Science and, Technology (JAMSTEC), Yokosuka, Japan; 3School of Geography, Earth and Environmental Sciences, Plymouth University, Plymouth, UK; 4National Oceanography Centre Southampton, University of Southampton, Southampton, UK; 5Department of Earth Sciences, Kanazawa University, Kanazawa, Japan *Corresponding author: E-mail: [email protected]
4.2.1.1 INTRODUCTION In March–April 1961, the drilling barge CUSS1 undertook the first scientific ocean drilling operation off Guadalupe Island, ∼240 km west of Baja California (Mexico). This expedition, beautifully reported in LIFE magazine by the novelist John Steinbeck and the renowned science photographer Fritz Goro, was the first (and eventually only) concrete manifestation of Project Mohole1. Project Mohole was an ambitious endeavor proposed initially in the late 1950s by the American Miscellaneous Society (AMSOC), an informal group of notable U.S. scientists, mostly geophysicists and oceanographers associated with the Office of Naval Research, including Harry Hess and Walter Munk. The principal aim was to drill through the oceanic crust, through the Mohorovičić Discontinuity, and to retrieve samples from Earth’s mantle (Teagle & Ildefonse, 2011). In his
1. http://www.nationalacademies.org/mohole.html. Developments in Marine Geology, Volume 7. http://dx.doi.org/10.1016/B978-0-444-62617-2.00017-7 Copyright © 2014 Elsevier B.V. All rights reserved.
449
450 Earth and Life Processes Discovered from Subseafloor Environments
book “A Hole in the Bottom of the Sea” (Bascom, 1961), Willard Bascom, Director of Project Mohole, notes that probably the first written suggestion for deep penetration down into the mantle was given by Frank Estabrook, an astrophysicist from the Basic Research branch of the U.S. Army in Pasadena (Estabrook, 1956). In addition to accomplishing a major technological breakthrough (including the invention of dynamic positioning, the drilling guide horn, and deepwater drill hole reentry), the Project Mohole drilling expedition in 1961 cored into oceanic basement for the first time, and demonstrated that the uppermost ocean crust consisted of basaltic lavas. This achievement received a personal letter of congratulations from U.S. President Kennedy2. More than 50 years later, and since the launch of DSDP in 1968, oceanic basement has been drilled in a range of geodynamic settings. The first intentional drilling efforts in the oceanic crust were DSDP Leg 34 in 1973–1974, on the Nazca Plate in the Eastern Pacific Ocean (Yeats et al., 1976), and DSDP Leg 37 in 1974 on the western flank of the Mid-Atlantic Ridge, south of the Azores Plateau (Aumento et al., 1977). These Legs recovered for the first time substantial cores of basalt from the upper oceanic crust. In the Atlantic, cores from Site 334 also recovered small amounts of gabbros and serpentinized peridotites for the first time at shallow subseafloor depths, demonstrating, together with future seafloor geology, marine geophysics, and drilling along the Mid-Atlantic Ridge, that the layered “Penrose” model from the early 1970s (Anonymous, 1972) is not applicable worldwide. Since then, scientific ocean drilling has significantly contributed to our understanding of the variability of the architecture of oceanic lithosphere. The style of accretion critically depends on the balance between magma production, hydrothermal cooling, and tectonics, which is on the first order related to spreading rate. Seismic, bathymetric, and marine geological observations indicate that ocean crust formed at fast spreading rates (full rate >80 mm per year) is much less variable than crust formed at slow spreading rates (<40 mm per year) and is closer to the ideal “Penrose” layered model developed from ophiolites (Anonymous, 1972). Yet, numerous basic, direct observations regarding the architecture of in situ, present-day ocean crust, including rock types, deformation, or geochemistry, are yet to be made. A compilation of holes into the ocean crust cored by scientific ocean drilling since the beginning of DSDP is presented in Table 4.2.1.1 and Figures 4.2.1.1 and 4.2.1.2. Only 37 holes deeper than 100 m have been cored in oceanic crust since DSDP Leg 37 in 1974 (Figure 4.2.1.2). The recovered material represents <2% of the cores recovered to date by DSDP, ODP, and IODP. In spite of this cursory sampling, scientific drilling has contributed significantly to advancing knowledge of ocean crust architecture, mid-ocean-ridge accretion, and hydrothermal processes (e.g., Alt et al., 1996; Blackman et al., 2011; Dick et al., 2000; Dick, Natland, & Ildefonse, 2006; Ildefonse et al., 2007a,b; Ildefonse, Rona, & Blackman, 2007c; Teagle, Alt, & Halliday, 1998; Wilson et al., 2006). 2. http://nationalacademies.org/includes/KennedyTelegram.pdf.
TABLE 4.2.1.1 Drill Holes into Oceanic Basement in Intact Crust and Tectonically Exposed Lower Crust and Upper Mantle. This Compilation Does Not Include Other “Hard Rock” Drill Holes in Oceanic Plateaus, Arc Basement, Hydrothermal Mounds, or Passive Margins. Age (my)
Sediment Thickness (m)
Basement Penetration (m)
Recovery (%)
Spreading Rate
Lithology
Comments
11°09.21′S 70°31.56′E
Indian
2844.5
30
506
81
50
S/I
Basaltic lavas
Projection of Chagos-Laccadive Plateau
257
30°59.16′S 108°20.99′E
Indian
5278
120
262
65
50
S/I
Basalt and breccia
Wharton Basin off Perth, Australia
34
319A
13°01.04′S 101°31.46′W
Pacific
4296
16
98
59
25
F
Basaltic lavas
Bauer Deep, 13°S East Pacific Rise
37
332A
36°52.72′N 33°38.46′W
Atlantic
1851
3.5
104
331
10
S
Basalt, basalt breccia, and interlayered sediments
Mid-Atlantic Ridge 36°–37°N
37
332B
36°52.72′N 33°38.46′W
Atlantic
1983
3.5
149
583
21
S
Basalt and Mid-Atlantic basalt breccia Ridge 36°–37°N
37
333A
36°50.45′N 33°40.05′W
Atlantic
1665.8
3.5
218
311
8
S
Basalt and Mid-Atlantic basalt breccia Ridge 36°–37°N
37
335
37°17.74′N 35°11.92′W
Atlantic
3188
15
454
108
38
S
Basaltic lavas
Mid-Atlantic Ridge 36°–37°N
45
395A
22°45.35′N 46°04.90′W
Atlantic
4485
7.3
92
571
18
S
Basaltic lavas and breccia
Mid-Atlantic Ridge 23°N
Hole
Location
24
238
26
Formation and Evolution of Oceanic Lithosphere Chapter | 4.2.1 451
Ocean
Water Depth (m)
Leg/ Expedition
Continued
TABLE 4.2.1.1 Drill Holes into Oceanic Basement in Intact Crust and Tectonically Exposed Lower Crust and Upper Mantle. This Compilation Does Not Include Other “Hard Rock” Drill Holes in Oceanic Plateaus, Arc Basement, Hydrothermal Mounds, or Passive Margins.—cont’d
Ocean
Water Depth (m)
Age (my)
Sediment Thickness (m)
Basement Penetration (m)
Recovery (%)
Spreading Rate
Lithology
Comments
22°58.88′N 43°30.95′W
Atlantic
4450
9
126
96
33
S
Basaltic lavas
Mid-Atlantic Ridge 23°N
396B
22°59.14′N 43°30.90′W
Atlantic
4459
13
151
255
23
S
Basalt and breccia
Mid-Atlantic Ridge 23°N
49
410A
45°30.53′N 29°28.56′W
Atlantic
2987
9
331
49
38
S
Basaltic lavas
Mid-Atlantic Ridge 45°N
49
412A
36°33.74′N 33°09.96′W
Atlantic
2626
1.6
163
131
18
S
Basalt flows and intercalating limestone
Mid-Atlantic Ridge 33°N
51–53
417A
25°06.63′N 68°02.48′W
Atlantic
5478.2
110
208
209
61
S
Basaltic lavas
Western Atlantic
51–53
417D
25°06.69′N 68°02.81′W
Atlantic
5489
110
343
366
70
S
Basaltic lavas
Western Atlantic
51–53
418A
25°02.10′N 68°03.44′W
Atlantic
5519
110
324
544
72
S
Basaltic lavas
Western Atlantic
54
428A
09°02.77′N 105°26.14′W
Pacific
3358.5
2.3
63
53
39
F
Basaltic lavas
9°N East Pacific Rise
63
469
32°37.00′N 120°32.90′W
Pacific
3802.5
17
391
63
34
I
Basaltic lavas
Off California Coast
Leg/ Expedition
Hole
Location
45
396
46
63
470A
28°54.46′N 117°31.11′W
Pacific
3554.5
15
167
49
33
I
Basaltic lavas
Off California Coast
65
482B
22°47.38′N 107°59.60′W
Pacific
3015
0.5
137
93
54
I
Massive basalt and interlayered sediment
Off Gulf of California
65
482D
22°47.31′N 107°59.51′W
Pacific
3015
0.5
138
50
50
I
Massive basalt and interlayered sediment
Off Gulf of California
65
483
22°53.00′N 108°44.90′W
Pacific
3084
2
110
95
40
I
Massive basalt and pillow basalt with interlayered sediments
Off Gulf of California
65
483B
22°52.99′N 108°44.84′W
Pacific
3084
2
110
157
47
I
Massive basalt and pillow basalt with interlayered sediments
Off Gulf of California
65
485A
22°44.92′N 107°54.23′W
Pacific
2996.5
1.2
153.5
178
51
I
Massive basalt and interlayered sediments
Off Gulf of California
Continued
TABLE 4.2.1.1 Drill Holes into Oceanic Basement in Intact Crust and Tectonically Exposed Lower Crust and Upper Mantle. This Compilation Does Not Include Other “Hard Rock” Drill Holes in Oceanic Plateaus, Arc Basement, Hydrothermal Mounds, or Passive Margins.—cont’d
Ocean
Water Depth (m)
Age (my)
Sediment Thickness (m)
Basement Penetration (m)
Recovery (%)
Spreading Rate
Lithology
Comments
1°13.63′N 83°44.06′W
Pacific
3466.9
6.6
264
73
60
I
Basaltic lavas
South Flank of Costa Rica Rift
504B
1°13.611′N 83°43.818′W
Pacific
3474
6.6
270
1841
20
I
Basalt, stockwork, and diabase
South Flank of Costa Rica Rift
82
559
35°07.45′N 40°55.00′W
Atlantic
3754
35
238
63
37
S
Basaltic lavas
West Flank of MidAtlantic Ridge 35°N
82
562
33°08.49′N 41°40.76′W
Atlantic
3172
12
240
90
45
S
Pillow basalt and massive basalt
West Flank of MidAtlantic Ridge 33°N
82
564
33°44.36′N 43°46.03′W
Atlantic
3820
35
284
81
43
S
Pillow basalt and minor massive basalt
West Flank of MidAtlantic Ridge 34°N
91
595B
23°49.34′S 165°31.61′W
Pacific
5615
80
70
54
28
F
Vesicular Central South Pacific aphyric basalt
92
597C
18°48.43′S 129°46.22′W
Pacific
4164
30
53
91
53
F
Massive basalt flows
Leg/ Expedition
Hole
Location
68
501
69/70/83/111/ 137/140/ 148
West Flank South East Pacific Rise 18°S
106/109
648B
22°55.32′N 44°56.825′W
Atlantic
3326
0
0
50
12
S
Pillow basalt
Mid-Atlantic Ridge 23°N
109
670A
23°9.996′N 45°1.932′W
Atlantic
3625
0
0
77
6
S
Serpentinized peridotite
Mid-Atlantic Ridge, MARK, 23°N
118/176
735B
32°43.395′S 57°15.959′E
Indian
720
11.8
0
1508
86
S
Gabbro
Atlantis Bank, Southwest Indian Ridge
123
765D
15°58.56′S 117°34.51′E
Indian
5713.8
140
928
267
31
F
North–east mid-oceanridge basaltic lavas
Argo Abyssal Plain
129/185
801C
18°38.538′N 156°21.59′E
Pacific
5674
170
462
414
47
F
Pillow basalt, basalt flows, and breccias
Western North Pacific
129
802A
12°5.778′N 153°12.63′E
Pacific
5980
120
509
51
33
F
Basaltic lavas
Western North Pacific
136
843B
19°20.54′N 159°5.68′W
Pacific
4418
95
243
71
37
F
Basaltic lavas
West of Hawaii
147
894E
2°18.059′N 101°31.524′W
Pacific
3014
1
0
29
11
F
Gabbro
Hess Deep
147
894F
2°17.976′N 101°31.554′W
Pacific
3025
1
0
26
7
F
Gabbro
Hess Deep
147
894G
2°17.976′N 101°31.554′W
Pacific
3023
1
0
127.5
35
F
Gabbro
Hess Deep
Continued
TABLE 4.2.1.1 Drill Holes into Oceanic Basement in Intact Crust and Tectonically Exposed Lower Crust and Upper Mantle. This Compilation Does Not Include Other “Hard Rock” Drill Holes in Oceanic Plateaus, Arc Basement, Hydrothermal Mounds, or Passive Margins.—cont’d
Ocean
Water Depth (m)
Age (my)
Sediment Thickness (m)
Basement Penetration (m)
Recovery (%)
Spreading Rate
2°16.638′N 101°26.766′W
Pacific
3821
1
0
17
14
895B
2°16.638′N 101°26.760′W
Pacific
3821
1
0
10
147
895C
2°16.632′N 101°26.772′W
Pacific
3820
1
0
147
895D
2°16.638′N 101°26.778′W
Pacific
3821
1
147
895E
2°16.788′N 101°26.790′W
Pacific
3753
147
895F
2°16.902′N 101°26.790′W
Pacific
148
896A
1°13.006′N 83°43.392′W
153
920B
153 153
Leg/ Expedition
Hole
Location
147
895A
147
Lithology
Comments
F
Serpentinized peridotite
Hess Deep
10
F
Serpentinized peridotite
Hess Deep
38
15
F
Serpentinized peridotite
Hess Deep
0
94
20
F
Serpentinized peridotite
Hess Deep
1
0
88
37
F
Serpentinized peridotite
Hess Deep
3693
1
0
26
8
F
Serpentinized peridotite
Hess Deep
Pacific
3459
6.6
179
290
27
I
Basaltic lavas
South Flank of Costa Rica Rift
23°20.310′N 45°1.038′W
Atlantic
3339
<1
0
126.4
35.3
S
Serpentinized peridotite
Mid-Atlantic Ridge, MARK, 23°N
920D
23°20.322′N 45°1.044′W
Atlantic
3338
<1
0
200.8
47.3
S
Serpentinized peridotite
Mid-Atlantic Ridge, MARK, 23°N
921A
23°32.460′N 45°1.866′W
Atlantic
2488
<1
0
17.1
18.1
S
Gabbro
Mid-Atlantic Ridge, MARK, 23°N
153
921B
23°32.478′N 45°1.842′W
Atlantic
2490
<1
0
44.1
19.4
S
Gabbro
Mid-Atlantic Ridge, MARK, 23°N
153
921C
23°32.472′N 45°1.830′W
Atlantic
2495
<1
0
53.4
11.4
S
Gabbro
Mid-Atlantic Ridge, MARK, 23°N
153
921D
23°32.442′N 45°1.830′W
Atlantic
2514
<1
0
48.6
12.7
S
Gabbro
Mid-Atlantic Ridge, MARK, 23°N
153
921E
23°32.328′N 45°1.878′W
Atlantic
2456
<1
0
82.6
21.4
S
Gabbro
Mid-Atlantic Ridge, MARK, 23°N
153
922A
23°33.162′N 45°1.926′W
Atlantic
2612
<1
0
14.6
63.2
S
Gabbro
Mid-Atlantic Ridge, MARK, 23°N
153
922B
23°31.368′N 45°1.926′W
Atlantic
2612
<1
0
37.4
25.6
S
Gabbro
Mid-Atlantic Ridge, MARK, 23°N
153
923A
23°32.556′N 45°1.896′W
Atlantic
2440
<1
0
70
57.2
S
Gabbro
Mid-Atlantic Ridge, MARK, 23°N
153
924B
23°32.460′N 45°0.858′W
Atlantic
3170
<1
0
30.8
8.7
S
Gabbro
Mid-Atlantic Ridge, MARK, 23°N
153
924C
23°32.496′N 45°0.864′W
Atlantic
3177
<1
0
48.5
23.1
S
Gabbro
Mid-Atlantic Ridge, MARK, 23°N
168
1025C
47°53.250′N 128°38.880′W
Pacific
2602
1.237
106
41
37
I
Basalt
Juan de Fuca Flank
168
1026B
47°45.759′N 127°45.552′W
Pacific
2658
3.511
256
39
5
I
Basalt
Juan de Fuca Flank
168
1026C
47°46.261′N 127°45.186′W
Pacific
2669
3.516
229
19
3.5
I
Basalt
Juan de Fuca Flank
Continued
TABLE 4.2.1.1 Drill Holes into Oceanic Basement in Intact Crust and Tectonically Exposed Lower Crust and Upper Mantle. This Compilation Does Not Include Other “Hard Rock” Drill Holes in Oceanic Plateaus, Arc Basement, Hydrothermal Mounds, or Passive Margins.—cont’d
Ocean
Water Depth (m)
Age (my)
Sediment Thickness (m)
Basement Penetration (m)
Recovery (%)
Spreading Rate
Lithology
Comments
47°46.776′N 128°7.320′W
Pacific
2645
2.621
290
48
6.5
I
Basalt
Juan de Fuca Flank
1105A
32°43.135′S 57°16.652′E
Indian
714
11.8
0
158
75
S
Gabbro
Atlantis Bank, Southwest Indian Ridge
185
1149D
31°18.79′N 143°24.03′E
Pacific
5818
133
307
133
17
F
Pillow basalt, basalt flows, and breccias
Western North Pacific
187
1162B
44°37.9′S 129°11.3′′E
Indian
5464
18
333
59
17
I
Basaltic lavas and breccia
Australian-Antarctic Discordance
187
1163A
44°25.5′S 126°54.5′E
Indian
4354
17
161
47
33
I
Basaltic lavas
Australian-Antarctic Discordance
187
1164B
43°45.0′S 127°44.8′E
Indian
4798
18.5
150
66
16
I
Basaltic lavas
Australian-Antarctic Discordance
191
1179D
41°04.8′N 159°57.8′E
Pacific
5563.9
129
377
98
44
F
Basaltic lavas
Western North Pacific
200
1224F
27°53.36′N 141°58.77′W
Pacific
4967.1
46
28
147
26
F
Basaltic lavas
Central Pacific
203
1243B
5°18.07′N 110°04.58′W
Pacific
3868
11
110
87
25
F
Basaltic lavas
Western Flank East Pacific Rise 5°N
Leg/ Expedition
Hole
Location
168
1032A
179
206
1256C
6°44.18′N 91°56.06′W
Pacific
3634.7
15
251
89
61
F
Basaltic lavas
Cocos plate eastern Flank East Pacific Rise
206/ 309/312
1256D
6°44.16′N 91°56.06′W
Pacific
3634.7
15
250
1257.1
37.1
F
Basaltic Cocos plate Eastern lavas, sheeted Flank East Pacific dike, and Rise varitextured gabbro
209
1268A
14°50.755′N 45°4.641′W
Atlantic
3007
0
147.6
53.3
S
Serpentinized peridotite
Mid-Atlantic Ridge 15°20′N
209
1270A
14°43.342′N 44°53.321′W
Atlantic
1951
0
26.9
12.2
S
Gabbro
Mid-Atlantic Ridge 15°20′N
209
1270B
14°43.265′N 44°53.225′W
Atlantic
1909
0
45.9
37.4
S
Gabbro
Mid-Atlantic Ridge 15°20′N
209
1270C
14°43.284′N 44°53.091′W
Atlantic
1822
0
18.6
10.6
S
Gabbro
Mid-Atlantic Ridge 15°20′N
209
1270D
14°43.270′N 44°53.084′W
Atlantic
1817
0
57.3
13.4
S
Gabbro
Mid-Atlantic Ridge 15°20′N
209
1271A
15°2.222′N 44°56.887′W
Atlantic
3612
0
44.8
12.9
S
Serpentinized peridotite
Mid-Atlantic Ridge 15°20′N
209
1271B
15°2.189′N 44°56.912′W
Atlantic
3585
0
103.8
15.3
S
Serpentinized peridotite
Mid-Atlantic Ridge 15°20′N
209
1272A
15°5.666′N 44°58.300′W
Atlantic
2560
0
131
28.6
S
Serpentinized peridotite
Mid-Atlantic Ridge 15°20′N Continued
TABLE 4.2.1.1 Drill Holes into Oceanic Basement in Intact Crust and Tectonically Exposed Lower Crust and Upper Mantle. This Compilation Does Not Include Other “Hard Rock” Drill Holes in Oceanic Plateaus, Arc Basement, Hydrothermal Mounds, or Passive Margins.—cont’d Leg/ Expedition
Hole
Location
Ocean
Water Depth (m)
Sediment Thickness (m)
Basement Penetration (m)
Recovery (%)
Spreading Rate
209
1274A
15°38.867′N 46°40.582′W
Atlantic
3940
0
155.8
22.2
209
1275B
15°44.486′N 46°54.208′W
Atlantic
1562
0
108.7
209
1275D
15°44.440′N 46°54.217′W
Atlantic
1554
0
301
U1301A 47°45.209′N 127°45.833′W
Pacific
2656
3.5
301
U1301B 47°45.229′N 127°45.826′W
Pacific
2655
304
U1309B
30°10.108′N 42°7.110′W
Atlantic
304/ 305
U1309D 30°10.120′N 42°7.113′W
Atlantic
Age (my)
Lithology
Comments
S
Serpentinized peridotite
Mid-Atlantic Ridge 15°20′N
43.1
S
Gabbro
Mid-Atlantic Ridge 15°20′N
209
50
S
Gabbro
Mid-Atlantic Ridge 15°20′N
262
108
0
I
Basaltic lavas
Juan de Fuca Ridge Flank; no coring (CORK)
3.5
265
318
12.9
I
Basaltic lavas
Juan de Fuca Ridge Flank; recovery is only for the 232 m of cored basement
1642
2
2
99.8
45.9
S
Gabbro
Mid-Atlantic Ridge 30°N
1645
2
2
1413.3
74.8
S
Gabbro
Mid-Atlantic Ridge 30°N
327
U1362A 47°45.663′N 127°45.672′W
Pacific
2661
3.5
236
292
29.6
I
Basaltic lavas
Juan de Fuca Ridge Flank; recovery is only for the 150 m of cored basement
336
U1383C 22°48.124′N 46°03.166′W
Atlantic
4414
8
38.3
293.2
26.5
S
Basaltic lavas
North Pond. Corked Hole. Upper 31.2 m of basement not cored
345
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S = Slow (<40 mm per year); I = Intermediate; F = Fast (>80 mm per year); MARK = Mid-Atlantic Ridge Kane Fracture Zone; CORK = Circulation Obviation Retrofit Kit.
462 Earth and Life Processes Discovered from Subseafloor Environments +ROHV'ULOOHGLQ2FHDQ&UXVW '6'3 2'3 ,2'3
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FIGURE 4.2.1.1 DSDP, ODP, and IODP holes drilled in ocean crust >100 mbsf from 1974 to 2013. Sites mentioned in the text are labeled. Also indicated are the locations of the four IODP expeditions reviewed herein. The seafloor age is based on age grid by Müller et al. (2008, revised version 3; www.earthbyte.org/). This map does not include “hard rock” drill holes in oceanic plateaus, arc basement, hydrothermal mounds, or passive margins.
Hole 504B, located on 6.9 my crust formed at an intermediate rate at the Costa Rica Rift (Figure 4.2.1.1), remains the second deepest hole (2111 meters below seafloor (mbsf)) in all of scientific ocean drilling and the deepest into oceanic basement (1836 m subbasement; Alt et al., 1996). This site was host to drilling and other experiments over eight DSDP and ODP cruises, and was the first hole to penetrate completely through the volcanic lava sequences and ∼1 km into sheeted dikes in intact, layered oceanic crust. It remains a reference hole for hydrothermal alteration of the upper ocean crust (e.g., Alt, Honnorez, Laverne, & Emmermann, 1986; Alt, Muehlenbachs, & Honnorez, 1986), and for the geological significance of seismic Layers 2A and 2B (e.g., Carlson, 2011), thought to represent a porous lava flow layer and the underlying sheeted dike complex, respectively (e.g., Herron, 1982). In fast-spread crust, the Layer 2/3 boundary is classically assumed to be the boundary between the upper crust (sheeted dikes and lavas) and the underlying plutonic gabbroic crust. This transition between typical layer 2 and layer 3 seismic velocity gradients with depth has been sampled in situ Hole 504B, where it lies in the sheeted dike complex, and appears related to changes in alteration and porosity structure (Carlson, 2010; Detrick, Collins, Stephen, & Swift, 1994).
Formation and Evolution of Oceanic Lithosphere Chapter | 4.2.1 463
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FIGURE 4.2.1.2 Compilation chart showing holes drilled >100 m in intact crust and tectonically exposed lower crust and upper mantle from 1974 to 2013 (drill hole locations in Figure 4.2.1.1). For each hole are indicated the hole number and the recovery (in percent) for each lithology. This compilation does not include “hard rock” drill holes in oceanic plateaus, arc basement, hydrothermal mounds, or passive margins.
A complete, continuous section of intact, layered fast-spread crust down to the cumulate gabbro layers has yet to be drilled and remains a first-order scientific target of ocean drilling for the ocean crust research community (e.g., Chikyu+10 Steering Committee, 2013; Dick & Mével, 1996; Ildefonse et al., 2010a; I ldefonse et al., 2010b; Ildefonse, Christie, & the Mission Moho Workshop S teering Committee, 2007b; IODP S cience Plan 2013–2023, 2011; Ravelo et al., 2010; Teagle, Wilson, & Acton, 2004; Teagle et al., 2009). Over the last 10 years, Hole 1256D (ODP Leg 206, Integrated Ocean Drilling Program (IODP) Expeditions 309, 312, and 335; Figure 4.2.1.1) penetrated to the base of the sheeted dike complex and the uppermost gabbro in superfast-spread crust (i.e., >130 mm per year; Lonsdale, 1977); this was the first sampling of the transition to plutonic rocks in intact ocean crust (Teagle et al., 2006; Teagle, Ildefonse, Blum, & the Expedition 335 Scientists, 2012; Wilson et al., 2003; Wilson et al., 2006). An alternative to drilling deep through intact, fast-spread crust is to access the deepest part of the crust in tectonic windows. This strategy was recently used at Hess Deep, in the eastern equatorial Pacific just off the East Pacific Rise (EPR), during IODP Expedition 345 to access lower crustal gabbroic rocks (Gillis, Snow, Klaus, & the Expedition 345 Scientists, 2014; Gillis et al., 2014a; Figure 4.2.1.1).
464 Earth and Life Processes Discovered from Subseafloor Environments
In 1986, serpentinized mantle peridotites were intentionally drilled at a midocean ridge for the first time during ODP Leg 109 (Site 670; Detrick et al., 1988) on the west wall of the Mid-Atlantic Ridge median valley near 23°10′N. In the same area, south of the Kane Fracture Zone, 95 m of serpentinized peridotites were recovered from a 200-m-deep hole during ODP Leg 153 (Site 920; Cannat et al., 1995). Ten years later, ODP Leg 209 (Sites 1268–1275) returned to drill in the peridotite-rich area around the 15°20′N fracture zone on the MidAtlantic Ridge (Kelemen, Kikawa, Miller, & Shipboard Scientific Party, 2007). Thirteen holes at six sites along the spreading axis penetrated mantle peridotite and gabbroic rocks in proportions that were roughly 70/30, similar to what was previously sampled by dredging and submersible on the seafloor in the same area (compiled data and references in Kelemen et al., 2004). These findings, together with geophysics and seafloor geology, contributed to the understanding that slowspreading oceanic crust can have a very variable architecture, locally consisting of relatively small gabbro bodies intruded into serpentinized mantle peridotites, and that spreading in nonlayered, magma-poor crust is largely accommodated by detachment faulting (e.g., Cannat, 1993; Dick, 1989; Escartín et al., 2008; Escartín & Canales, 2011; Kelemen et al., 2007; Lagabrielle, Bideau, Cannat, Karson, & Mével, 1998). In this type of crust, the spatially and temporally variable interplay of magmatic and tectonic processes is locally expressed in the development of Oceanic Core Complexes (OCCs). OCCs are domal features that result from the exhumation of lower crustal and upper mantle rocks through long-lived (typically 1–3 my) detachment faulting (e.g., Cann et al., 1997; Dick, Tivey, & Tucholke, 2008; Escartín & Canales, 2011; Escartín, Mével, MacLeod, & McCaig, 2003; MacLeod et al., 2009; Smith, Escartín, Schouten, & Cann, 2008; Tucholke & Lin, 1994; Tucholke, Lin, & Kleinrock, 1998). The first deep hole drilled in an OCC was the 1508-m-deep ODP Hole 735B in gabbroic rocks of the Atlantis Bank on the Southwest Indian Ridge (Dick et al., 2000). More recently, IODP successfully sampled another OCC, with Hole U1309D penetrating 1415 m of gabbroic rocks at the Atlantis Massif on the Mid-Atlantic Ridge during Expeditions 304 and 305 (Blackman et al., 2011; Ildefonse et al., 2007a). This chapter draws from results of 10 years of IODP ocean drilling (2003– 2013) in igneous oceanic crust, focusing on the two deep holes in slow-spread (Hole U1309D) and fast-spread (Hole 1256D) crust, and briefly summarizing the outcome of the recent IODP Expedition 345 at Hess Deep. These drilling expeditions have improved our understanding of crustal architecture and accretion processes at mid-ocean ridges.
4.2.1.2 DEEP DRILLING IN SLOW-SPREAD CRUST: THE ATLANTIS MASSIF 4.2.1.2.1 Background IODP Expeditions 304 and 305, from November 2004 to March 2005, targeted one of the emblematic OCCs, the Atlantis Massif, at 30°N on the Mid-Atlantic
Formation and Evolution of Oceanic Lithosphere Chapter | 4.2.1 465
Ridge (Blackman, Cann, Janssen, & Smith, 1998; Blackman et al., 2002; Cann et al., 1997; Schroeder & John, 2004). The Atlantis Massif is a young (0.5–2 my) OCC, situated at the inside corner of the intersection between the Mid-Atlantic Ridge and the Atlantis Fracture zone at 30°N. Seafloor lithologies observed by dredging and submersible sampling include serpentinites, gabbros, and basaltic rocks (Figure 4.2.1.3; Cann et al., 1997; Blackman et al., 2002). The active, serpentinite-hosted Lost City hydrothermal vent field (Früh-Green et al., 2003; Kelley et al., 2001) is located immediately south of the summit of the massif. Localized deformation and fluid flow associated with detachment faulting are documented from seafloor sampling on the southern wall of the massif (Boschi, Früh-Green, Delacour, Karson, & Kelley, 2006; Karson et al., 2006; Schroeder & John, 2004). Two deep holes were drilled at Site U1309 during IODP Expeditions 304 and 305, in the footwall of the detachment fault (Figure 4.2.1.3; Blackman et al., 2006, 2011). In Û :
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466 Earth and Life Processes Discovered from Subseafloor Environments
addition, IODP Expedition 340T revisited this site in 2012 to conduct 5 days of borehole logging in Hole U1309D, to complete the logging program that was hampered by bad weather conditions at the end of IODP expedition 305 (Blackman et al., 2006). The program included completing sonic logging to the bottom of the hole, and acquiring new seismic imaging data from a Vertical Seismic Profile (VSP) experiment (Expedition 340T Scientists, 2012).
4.2.1.2.2 Revisiting the Geophysical Signature of the Atlantis Massif Based on the dominant occurrence of serpentinized mantle peridotites sampled on the south flank of the southern ridge (Figure 4.2.1.3; e.g., Blackman et al., 2002; Boschi et al., 2006), and on the predrilling interpretation of available seismic (Figure 4.2.1.3; Collins & Detrick, 1998; Collins, Tucholke, & Canales, 2002; Canales, Tucholke, & Collins, 2004) and gravimetric (Blackman, Cann, Janssen, & Smith, 1998; Nooner, Sasagawa, Blackman, & Zumberge, 2003) data, one of the goals of IODP Expeditions 304/305 was to drill and sample the core of an OCC that was inferred to be mostly made of mantle peridotites. It was hoped that fresh mantle peridotite was present at reasonably shallow depths (∼800 mbsf) accessible by conventional nonriser drilling and hence, could be sampled for the first time (see Blackman et al., 2011; for a summary of the predrilling rationale). The recovery of a 1415-m-long gabbroic section demonstrated that geological sampling by drilling is essential to better constrain the interpretation of geophysical data in ocean crust, particularly in slow-spread crust, which can be lithologically heterogeneous on small scales. Drill cores, together with wireline geophysical measurements provided critical observations for reexamining the precruise geophysical interpretations and for proposing a revised analysis of the lithological architecture of the Atlantis Massif and other OCCs (Blackman, Karner, & Searle, 2008; Blackman, Canales, & Harding, 2009; Blackman and Collins, 2010; Canales, Tucholke, Xu, Collins, & DuBois, 2008; Collins, Blackman, Harris, & Carlson, 2009; Henig, Blackman, Harding, Canales, & Kent, 2012). The positive 30–40 mGal residual gravity anomaly centered on the core of the OCC is now considered to reflect a gabbroic core with a density of 2900 kg/m3, the adjacent basaltic block is significantly fractured with an average density of 2600 kg/m3, and the portion of the lithosphere deeper than 1500 mbsf that has density lower than mantle rock (gabbroic, and/or significantly altered peridotite) is ∼3-km thick within the domal core (Blackman, Karner, & Searle, 2008). Postdrilling seafloor refraction modeling showed that shallow mantle velocities are not required (Figure 4.2.1.4; Collins et al., 2009) and velocities in the upper several hundred meters are typical of mafic rock (≤6.5 km/s). Tomographic inversion of refracted arrivals recorded by a 6-km multichannel seismic (MCS) streamer (Canales et al., 2008) produced a similar velocity structure to 1.0–1.5 km depth across the Central Dome (Figure 4.2.1.4(D)). Further detailed processing of the MCS data, continued downward using the “synthetic ocean bottom experiment” technique (Henig et al., 2012),
Formation and Evolution of Oceanic Lithosphere Chapter | 4.2.1 467
2
0.3
4
5
6
7
South
North
Southern Ridge
Central Dome
7
South
North
6 5 4
3
Hole U1309D
-15
(C) Depth (km)
3
0
-10
3
-5
0
5
2
-15
Hole U1309D
-10
2
8 7 6 5 4
1
3
10 km
-5
0
5 km
3 2
velocity (km/s)
Depth (km)
(B) 1
velocity (km/s) 2
velocity (km/s)
Depth (km)
(A)
Two-way travel time (s)
(D)
CMP FIGURE 4.2.1.4 Seismic imaging of the Atlantis Massif. (A) Preferred velocity model along the NOBEL refraction Line 9 on the central dome (location in Figure 4.2.1.3), determined from forward modeling by Collins et al. (2009). (B) Velocity model based on inversion of multichannel seismic (MCS) refractions for Line 4 (location in Figure 4.2.1.3) by Canales et al. (2008). (C) Velocity model based on inversion of MCS refractions for Line 4 by Henig et al. (2012). (D) Time-migrated section for MCS Line 4 (Canales et al., 2004); D: D-reflector; CMP: Common-Midpoint.
reveals significant spatial (both along-axis and along-flow line) variability of the P-wave velocity structure at the scale of the Atlantis Massif. This variability is generally consistent with lithologically heterogeneous crust, made of gabbroic plutons intruded into a dominantly peridotitic (serpentinized) crust. The western two-thirds of the southern ridge, to date unsampled by drilling, is dominated by
468 Earth and Life Processes Discovered from Subseafloor Environments
low-velocity material that is likely to be serpentinized peridotites (at least down to ∼1 km below seafloor), whereas the eastern part displays a higher velocity body that is interpreted as being a gabbro pluton. Analysis of a 40-km-long EW air gun refraction profile across the central dome provided constraints on velocities to depths as great as 7 km (Blackman and Collins, 2010). Traveltime tomography suggests that significant volumes of rock with mantle-like velocity (>7.5 km/s) only occur more than ∼5 km below seafloor within the dome, and only below 6 km in the axial valley. Initial results from IODP Expedition 340T seem to indicate that combined lithological changes and local increases in alteration could produce an impedance contrast strong enough to account for the reflectivity in the upper part of the section. In particular, the sonic log confirms that the borehole velocity of altered olivine-rich troctolite intervals in Hole U1309D may be sufficiently distinct from that of surrounding rocks (P-wave velocity (VP) ∼0.5 km/s slower) to produce an MCS reflection (Expedition 340T Scientists, 2012). However, the source(s) of the impedance contrast that gives rise to the strong reflection imaged throughout much of the dome (the “D-reflector”; Canales et al., 2004) is yet to be identified.
4.2.1.2.3 The Interplay between Magmatism and Tectonics Controls the Development of OCCs Hole U1309D sampled a 1415-m-long section of gabbroic rocks in the core of the Atlantis Massif OCC, with 75% recovery (Figure 4.2.1.5). This is the second longest core drilled in slow-spread crust, after ODP Hole 735B in the Southwest Indian Ocean (Dick et al., 2000). The borehole lithology at Site U1309D contrasts remarkably with the seafloor lithology of the Atlantis M assif. Although seafloor sampling dominantly recovered serpentinized peridotites on the south face of the Southern Ridge (approximately 70% of the recovered samples; Blackman et al., 2002; Karson et al., 2006; Schroeder & John, 2004), only three thin (<1 m) intervals of ultramafic rocks interpreted as residual mantle peridotites, intercalated within gabbros in the upper 225 m of the section (Tamura, Arai, Ishimaru, & Andal, 2008), were recovered in Holes U1309B and U1309D on the Central Dome, 5 km to the north. It is worth noting that including the completion of IODP Expeditions 304/305, ODP and IODP has drilled 16 holes in four different OCCs, and that all holes had recovered exclusively gabbroic sections (Ildefonse et al., 2007a). Hence OCCs are associated with significant magmatism. In contrast with early models (e.g., Escartin et al., 2003; Tucholke & Lin, 1994; Tucholke et al., 1998), this basic observation indicates that OCCs are not the magma-starved endmember of slow-spreading ridges. The revised model proposed that, in heterogeneous crust composed of gabbroic plutons intruded into variably serpentinized mantle peridotites (resulting from episodic magmatic activity), the development of an OCC is associated with, or results from a period of enhanced magmatism at depth. If this model is correct, the
(A)
(B)
0
0.5
1
(C)
Mg# (= Mg/[Mg + Fe]) (whole rock [molar])
20 m running average Fraction of cored interval
50 0
60
70
80
90
(D)
Plastic and magmatic foliation (intensity, 5 m ave.)
Cataclasis and veining (intensity, 5 m ave.)
0
0
1
2
2
4
(E)
Total alteration (%, 5 m intervals) 0
20
40
60
(F)
(G)
Bulk density (g/cc)
80 2.4 2.6 2.8
3
3.2
(H)
Vp (Km/s) 4
5
6
(I)
High-stability magnetization (inclination [°]) 7
-60
-30
0
Zircon Pb/U Age (my) 1.0
1.1
1.2
1.3
1.4
1.17 ± 0.02 my
400
800
my 1.24 ± 0.02 my
Depth (mbsf)
600
1000
1200
1400 Basalt/Diabase Oxide gabbro Gabbro, gabbronorite Troctolitic gabbro, olivine gabbro Troctolite Olivine-rich troctolite, dunite, harzburgite
Basalt/Diabase Oxide gabbro Gabbro, gabbronorite, olivine-bearing gabbro Olivine gabbro, troctolitic gabbro
Magmatic Plastic
Veining Cataclasis
Discrete samples Logging runs 1 & 3 Logging runs 1 & 3 5 m running average
Discrete samples Logging: 304 5m running 305 average 340T
Single measurement Multiple measurements Discrete samples
Felsic dike Oxide gabbro
Geocentric axial dipole
Troctolite Olivine-rich troctolite, dunite, harzburgite
my Fault zones inferred from cataclastic deformation of fault gouge in core
Leucocratic gabbro
FIGURE 4.2.1.5 Downhole plots of data for IODP Hole U1309D (modified from Ildefonse et al., 2006). (A) Lithology (20 m running average, white = no recovery); (B) Mg# (shipboard data); (C) plastic and magmatic foliation intensity (shipboard macroscopic description); (D) cataclasis and veining intensity (shipboard macroscopic description); (E) total alteration (shipboard macroscopic description); (F) bulk density (from borehole and discrete sample shipboard measurements); (G) compressional velocity (from borehole and discrete sample shipboard measurements); (H) inclination of stable remanent magnetization (shipboard measurements); (I) ages obtained for core samples from zircon (postcruise data, Grimes et al., 2008. The gray vertical bars are interpretative coherent age blocks, above and below 600 mbsf, with the average ages indicated aside).
Formation and Evolution of Oceanic Lithosphere Chapter | 4.2.1 469
200
470 Earth and Life Processes Discovered from Subseafloor Environments
domal morphology of OCCs would essentially be the consequence of having larger gabbroic plutons being exhumed and unroofed by the associated detachment fault (Blackman et al., 2011; Ildefonse et al., 2007a). The model implies that serpentinized fault zones envelop gabbro bodies, explaining why serpentinites can be the dominant lithology on the top or flanks of OCCs, while drilling to date suggests that their core is dominantly gabbroic. The inference that OCCs are associated with significant magma emplacement in the crust is consistent with a series of recent numerical models, which suggest that the optimal magmatic input for the development of long-living detachment faults and OCCs corresponds to about half of the total spreading (Behn & Ito, 2008; Buck, Lavier, & Poliakov, 2005; Olive, Behn, & Tucholke, 2010; Tucholke, Behn, Buck, & Lin, 2008). It is also consistent with observations from the southwest Indian Ocean (Cannat et al., 2006), where the true amagmatic stage of spreading at the Southwest Indian Ridge corresponds to the development of a smooth topography through conjugate, long-lived detachment faulting, marked by broad ridges and the absence of fault scarps and volcanics. OCCs there are identified through their corrugated topography, and are interpreted to represent the intermediate level of magmatic input between the fully magmatic volcanic seafloor and the amagmatic smooth seafloor.
4.2.1.2.4 Paleomagnetic Data and Borehole Imaging Constrain Tectonic Rotation in OCCs Competing models for the development of OCCs through long-lived detachment faulting involve either displacement on planar, low-angle normal faults (e.g., Wernicke, 1995) or progressive shallowing by rotation of initially steeply dipping (>45°) faults as a result of flexural unloading (i.e., the “rolling-hinge” model; e.g., Lavier, Buck, & Poliakov, 1999). Magmatic rocks acquire a magnetization in the direction of the ancient geomagnetic field at the time of their cooling and crystallization that may be used as a marker for subsequent tectonic rotation. The different models of OCC development can be tested by comparing the orientation of paleomagnetic remanence vectors in recovered magmatic rocks to the expected orientation of the Earth’s magnetic field at the sampling location. However, a fundamental difficulty inherent in scientific ocean drilling is that core samples lack azimuthal orientation in the geographical reference frame. Hence, documenting the amount and style of tectonic rotation of OCC footwall sections using paleomagnetic data becomes a problem with too many unknowns and multiple solutions. The way around this problem is to combine paleomagnetic data with structural observations in the core and borehole wireline imaging (e.g., Haggas, Brewer, Harvey, & MacLeod, 2005; MacLeod, Parson, Sager, & the ODP Leg 135 Scientific Party, 1992). In the absence of borehole imaging, Garcés and Gee (2007) had to make an assumption about the orientation of the rotation axis for the proposed tectonic rotations inferred from paleomagnetic measurements in serpentinite and gabbro sections exhumed by
Formation and Evolution of Oceanic Lithosphere Chapter | 4.2.1 471
detachment faulting in the 15°20′N fracture zone region of the Mid-Atlantic Ridge, sampled during ODP Leg 209 (Kelemen et al., 2007). For the first time, Hole U1309D offered complete, high-quality Formation MicroScanner (FMS) imaging of the borehole coupled with high core recovery, enabling a comprehensive core-log integration study (Figure 4.2.1.6). Using paleomagnetic remanence data and FMS images to reconstruct the true orientation of core pieces, Morris et al. (2009) demonstrated that the upper 400 m of the footwall of the detachment fault in the Atlantis Massif experienced an ∼45° anticlockwise rotation around a ridge-parallel horizontal axis trending NNE (010°), providing unequivocal validation of the rolling-hinge model (e.g., Lavier et al., 1999) for the development of OCCs. Analysis of the high-quality FMS imagery from Hole U1309D also allowed the distribution of structures within the Atlantis Massif OCC footwall to be comprehensively documented (Pressling, Morris, John, & MacLeod, 2012), revealing details of the internal structure of an OCC
IODP core “X”
(B)
True N
Discrete samples
(A)
(C)
011° ± 6°
N
46° ± 6° CCW
Archive half core
Mean remanence (discrete samples) Reference vector
98 - 367 mbsf IODP core coordinates
True geographic coordinates
FIGURE 4.2.1.6 Effect of FMS-based core reorientation on paleomagnetic data from Hole U1309D, and resulting constraints on footwall rotation in the Atlantis Massif Oceanic Core Complex (OCC) (modified after Morris et al., 2009). (A) Paleomagnetic remanence directions from discrete samples (red symbols (dark gray in print versions)) and archive half core sections (green symbols (light gray in print versions)) in the unoriented IODP core reference frame have a scattered distribution as a result of the drilling process. (B) The same data after individual core pieces have been reoriented to a true geographic reference frame by matching structures observed in the core with the corresponding structures in the borehole wall (the orientation of which may be determined from the FMS imagery). The independently reoriented paleomagnetic data now display a clustered distribution from which a reliable mean remanence vector may be calculated. (C) Comparison of the full mean remanence vector from discrete samples with the reference geocentric axial dipole vector for Atlantis Massif allows the orientation of the axis of footwall rotation (tilting) and angle of rotation to be determined, indicating substantial rotation around a Mid-Atlantic Ridge (MAR)-parallel axis consistent with rolling-hinge models for OCC development.
472 Earth and Life Processes Discovered from Subseafloor Environments
for the first time. A zone of brittle fracturing dominates the upper 385 m of the hole, and structures indicating detachment-related damage zones extend down to 750 mbsf into the OCC footwall (See also Michibayashi et al., 2008).
4.2.1.2.5 Protracted Construction of the Lower Crust Cores recovered in the upper 120 m at Site U1309 contain 35–45% of diabase/basalt in the form of subhorizontal, 5- to 10-m-thick intrusive bodies (Blackman et al., 2011) with chilled margins against cataclastic gabbros, amphibole-rich breccias, and undeformed diabase. They are interpreted as dikes intruded into the detachment fault prior to tectonic rotation (McCaig & Harris, 2012). The composition of all diabase recovered at Site U1309D falls within the range of basalt compositions for the Mid-Atlantic Ridge in the 30°N region (Godard et al., 2009). The dominantly gabbroic sequence in Hole U1309D comprises hundreds of individual lithologic units ranging in composition from oxide-gabbro to olivine-rich troctolite (Figures 4.2.1.5 and 4.2.1.7), identified on the basis of modal composition, grain-size changes, and/or textural changes. Units vary in thickness, from centimeters to tens of meters (John et al., 2009). Many intrusive contacts were recovered and indicate a general relationship of more evolved rock types intruded into more primitive rock types, although the converse sense of intrusion is also observed. Leucocratic intrusions (centimeter-scale thickness) cut the sequence and, together with oxide-gabbro intervals, host zircon grains that have been used to obtain crystallization ages of the recovered crustal section. Ages range from 1.08 ± 0.07 my to 1.28 ± 0.05 my, which indicates that the magmatic accretion of the multiple units described above has occurred over a minimum of approximately 100–200 ky (Figure 4.2.1.5; Grimes, John, Cheadle, & Wooden, 2008). The distribution of ages is inconsistent with models requiring constant depth melt intrusion beneath the detachment fault. Instead, data suggest a protracted construction model (Grimes et al., 2008), in which gabbroic rocks intrude at various depths below the ridge axis over a depth range of at least 1.4 km, before being exhumed by the detachment fault. Modal mineralogy and bulk rock geochemistry (e.g., Figure 4.2.1.7(A) and (B)) of the gabbroic sequence are typical of a cumulate series crystallizing from a mid-ocean ridge basalt (MORB) source. However, the bulk composition of Hole U1309D does not represent the complement to basalts recently erupted at the nearby spreading axis (Godard et al., 2009). The composition of the recovered rocks spans a large spectrum (Figure 4.2.1.7), with some troctolites and gabbros, respectively having some of the highest and lowest MgO contents of oceanic crustal rocks sampled to date. The bulk trace element and major element chemistry are consistent with most or all of the section being crystallized from a MORB melt at depth. When all lithologies, including leucocratic intervals, are accounted for, an extra basaltic counterpart to this section is not required (Figure 4.2.1.8; see detailed discussion in Godard et al., 2009).
Formation and Evolution of Oceanic Lithosphere Chapter | 4.2.1 473
(C)
(A) Ol-rich troctolite Troctolite
Gabbro
Gabbro and Ol-gabbro Other Exp 304-305 samples
Intrusive contact Troctolite
(B)
Obscured contact
Reaction zone
Oxide gabbro
(D) Ol-rich troctolite
Atlantis massif (Southern wall)
Troctolite
Gabbro (SWIR)
Olivine gabbro
Gabbro (MAR)
Gabbro
Gabbro & troctolite (MAR-ODP Site 1275)
Gabbronorite Oxide gabbro
Basalt (MAR)
Leucocratic dike
Peridotite (MAR-15°20’N)
Basalt Diabase Peridotite
Peridotite (MARK) Impregnated Peridotite (MAR-15°20’N)
Oxide gabbro
Gabbro 2 cm
FIGURE 4.2.1.7 Petrology, geochemistry, and photographs of Site U1309 core samples (From Blackman et al., 2011). (A) Compositions of representative suite of samples from site U1309 are shown by small gray symbols. Other symbols indicate samples from within the intervals that are dominantly olivine-rich troctolite that were studied in detail by Drouin et al. (2009). (B) Bulk rock chemistry for Site U1309 and comparison with other areas (Godard et al., 2009). Note how troctolites and olivine-rich troctolites fill in previously sparsely populated Ni-Mg# space. (C) Photo of Core 304-U1309D-69R-1 shows intrusive contact between gabbro and troctolite and lower contact that is a reaction zone between the troctolite and a later oxide gabbro injection. (D) Oxide gabbro dike intrudes gabbro (Hole U1309D, 386 mbsf) with a sharp lower contact with gabbro (Grimes et al., 2008).
474 Earth and Life Processes Discovered from Subseafloor Environments
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FIGURE 4.2.1.8 Bulk composition of Hole U1309D compared to the compositions of bulk gabbroic crust at Hole 735B and of primary mid-ocean ridge basalt (MORB) magma illustrated on (A) Ca# versus Mg# and (B) extended trace element diagrams (modified from Godard et al., 2009). The bulk composition of Hole U1309D is calculated taking into account all gabbroic rocks (BH) and without the olivine-rich troctolites (BG). The estimate of the bulk composition of Hole U1309 (BH) is characterized by high Mg# compared to primary magma compositions. These high Mg# may indicate a cumulate composition complementary to that of MORB, but also an overestimation of the mass of the most primitive endmember of the gabbroic suite, or an underestimation of the mass of its most evolved endmembers in our calculations. These hypotheses are tested using two models: (1) BH + LD: a 95/5% mix of BH and leucocratic dyke (LD); (2) BG + LD: a 95/5% mix of BG and LD. Estimates of the bulk composition of Hole 735B (Hart et al., 1999; Dick et al., 2000) as well as the mean compositions of basalts at Site U1309 and Hole 735B are shown for comparison. Hole 735B “protolith estimate” was determined on the basis of analyses conducted on a small suite of fresh and unaltered samples representative of the “characteristic lithologies”
Formation and Evolution of Oceanic Lithosphere Chapter | 4.2.1 475
Relatively low-pressure crystallization depths (≤200 MPa, ∼6 km) are inferred based on the modal relationships and chemistry of cumulus and intercumulus phases (Godard et al., 2009; Suhr, Hellebrand, Johnson, & Brunelli, 2008). This conclusion is supported by a combination of rock ages, assumptions about fault geometry, and isotherm depth that constrain crystallization depths at 5–7 km (Grimes, Cheadle, John, Reiners, & Wooden, 2011).
4.2.1.2.6 Olivine-Rich Troctolites: Mantle Contribution to the Igneous Lower Crust Olivine-rich troctolites (>70% olivine) are present as relatively thin (approximately 1–12 m) units within two main intervals in Hole U1309D (approximately 310–350 mbsf and 1090–1235 mbsf; Figure 4.2.1.5). Bulk rock geochemical signatures of these troctolites are more primitive (high MgO) than other mafic samples from the ocean crust (Figure 4.2.1.7; Drouin, Godard, Ildefonse, Bruguier, & Garrido, 2009; Godard et al., 2009). They display classical cumulate textures (Figure 4.2.1.9; Blackman et al., 2006). However, several lines of evidence suggest that these rocks are not simply the first crystallized product of a closed, fractionating magma body. Trace element patterns calculated for melts in equilibrium with measured in situ plagioclase, clinopyroxene, and olivine compositions (Figure 4.2.1.9(A)) illustrate that the olivine is in disequilibrium with the MORB melts that crystallized the plagioclase and clinopyroxene (Drouin et al., 2009). Mineral chemistry, textural relations, and crystallographic preferred orientations indicate that the olivine-rich troctolites are the end product of melt–rock interaction between an initial olivine-bearing rock and MORB melt (Drouin, Godard, Ildefonse, Bruguier, & Garrido, 2009; Drouin, Ildefonse, & Godard, 2010). The observed, relatively stronger concentrations of olivine [001] axes in some of these rocks are interpreted as resulting from melt– rock interaction. A first stage of dunitization weakens the olivine crystallographic preferred orientation (e.g., Tommasi, Godard, Coromina, Dautria, & Barsczus, 2004). It is followed by the disaggregation of mantle peridotite
of Hole 735B; the “strip average” estimate was obtained using a strip-sampling technique (see Hart et al., 1999) to take into account alteration, small-scale lithologic variability, and patchiness of veins, and thus obtain a possibly more representative sampling. The gray field indicates values of primary MORB melts which represent the most primitive basaltic compositions expected at a spreading centre (data from Kinzler and Grove (1993) in (A) and from Workman and Hart (2005) in (B)). Back stripping the contribution of olivine at Hole U1309 gives a composition for the bulk gabbroic crust at Hole U1309 consistent with estimates of the Mg#–Ca# composition of primary MORB magmas (BG). However, this estimate is characterized by trace element signatures indicative of plagioclase accumulation. Only models taking into account the contribution of leucocratic dykes in the bulk hole estimates give results consistent with primary magma compositions for both major and trace elements.
476 Earth and Life Processes Discovered from Subseafloor Environments
(A)
?
(B)
(C)
FIGURE 4.2.1.9 Sample characteristics from olivine-rich troctolite intervals (modified from Blackman et al., 2011). (A) Computed rare earth element compositions of melts in equilibrium with the measured mineral chemistry (Drouin et al., 2009). (B) Rounded and subhedral olivine crystals (Ol) included in large clinopyroxene (Cpx) and plagioclase poikiloblasts forming poikilitic texture in olivine-rich troctolite (sample 345_U1309D_247R3_16–18). (C) Coarse‐grained subhedral olivine crystals with smooth edges and well-defined and widely spaced subgrain boundaries and poikiloblastic plagioclase (Pl) and orthopyroxene (Opx) (sample 345_U1309D_248R2_22–24). (B) and (C) are cross-polarized photomicrographs from Drouin et al. (2010).
(Figure 4.2.1.10), which, if not heavily fluxed with melt, would be expected to display [100] preferred alignment related to crystal–plastic deformation (Drouin et al., 2010). Disruption of preexisting high-temperature crystallographic preferred orientations during melt influx is also supported by the common occurrence of adjacent grains with similar crystallographic orientations (Drouin et al., 2010; Suhr et al., 2008). Although not unequivocally demonstrated, a mantle origin for the olivine-rich troctolites is favored (Figure 4.2.1.10; Drouin et al., 2009, 2010; Suhr et al., 2008). This proposed scenario suggests a new mechanism for the formation of the lowermost oceanic crust in which small volumes (approximately 1–10-m-thick intervals in the core) of mantle peridotite are impregnated by, and react with, basaltic melts until they become isolated and incorporated into the crust during protracted periods of construction by igneous intrusion. Whatever the initial lithology, impregnation by large volumes of melt has strongly modified the original composition and microstructure of the host rock, and a melt/rock ratio of 3:1 is estimated, based on a fractionation model to match the Mg, Ni, and Cr contents measured in olivine grains (Suhr et al., 2008). The few recovered intervals of mantle peridotites (see above) also show petrographic and geochemical evidence of melt–rock reactions (Tamura et al., 2008; Godard et al., 2009; von der Handt & Hellebrand, 2010).
Formation and Evolution of Oceanic Lithosphere Chapter | 4.2.1 477
MORBtype melt
Ol-rich troctolites [100]
ridge axis
lithosphere
[010]
[001]
Ol
Impregnation : - Melt fraction ≥ ~30% - Disruption of olivine framework & modification of CPO
2
Dunite Ol
[100]
[010]
[001]
asthenosphere
Dunitization : Weakening of original CPO
1
[100]
[010]
[001]
Ol
Opx
Harzburgite FIGURE 4.2.1.10 Sketch of the proposed scenario for the formation of Hole U1309D olivine-rich troctolites (Drouin et al., 2010). Percolation of basaltic melts through depleted mantle peridotite at near-solidus temperature results in dunitization (orthopyroxene dissolution, precipitation of olivine, and formation of dunite), with weakening of olivine crystallographic preferred orientation (CPO) (indicated by 1), and impregnation when increasing melt fraction (≥30%) locally disrupts the olivine framework and further modifies the olivine CPO (indicated by 2). The dunitization and impregnation are viewed as successive stages of a continuous process. As temperature decreases, melt crystallizes plagioclase and clinopyroxene within the assimilated olivine matrix to form o livine‐rich troctolite. The schematic cross-section of the ridge axis is modified from Cannat (1996). The examples of original harzburgite and dunite CPOs are from Polynesian xenoliths (RPA18A and UAH280X, respectively; Tommasi et al., 2004). Note the absence of structural reference frame (no visible foliation in samples) for the dunite and olivine-rich troctolite CPOs.
4.2.1.3 DEEP DRILLING OF INTACT OCEAN CRUST FORMED AT A SUPERFAST SPREADING RATE: HOLE 1256D 4.2.1.3.1 Background The vast majority (∼70%) of magma derived from the mantle is brought into the Earth’s crust at mid-ocean ridges, and approximately two-thirds of that magma cools and crystallizes in the lower portion of the oceanic crust. Although <20% of modern ridges are spreading at fast rates, nearly 50% of present-day ocean crust (i.e., ∼30% of the Earth’s surface) was formed at fast-spreading ridges. Because fast-spread crust is more continuous and regularly layered than slowspread crust, extrapolating fast-spreading accretion processes from a few sites
478 Earth and Life Processes Discovered from Subseafloor Environments
might reasonably describe a significant portion of the Earth’s surface. Additionally, the great majority of crust subducting into the mantle over the past ∼200 my formed at fast-spreading ridges (e.g., Müller et al., 2008), making characterizing this style of crust most relevant for understanding the recycling of crustaland ocean-derived components back into the mantle. Drilling a complete in situ section of ocean crust has been an unfulfilled ambition of Earth scientists for many decades, which provided the impetus for the conception of scientific deep ocean drilling (e.g., Teagle & Ildefonse, 2011); ODP Hole 1256D is the latest, and to date most successful attempt to reach this goal. The deep drilling campaign to Site 1256 was designed to understand the formation, architecture, and evolution of ocean crust formed at fast spreading rates, and has been the focus of four scientific ocean drilling cruises (ODP 206, Wilson et al., 2003; IODP Expedition 309/312, Teagle et al., 2006; IODP Expedition 335, Teagle et al., 2012). Hole 1256D is located in 3635 m of water in the Guatemala Basin (6°44.2′N, 91°56.1′W), on the Cocos plate in the eastern equatorial Pacific Ocean. Ocean crust here formed 15 my ago during a sustained episode of superfast ocean ridge spreading (>200 mm per year; Wilson, 1996; Figures 4.2.1.1 and 4.2.1.11) at the EPR. The site formed on a ridge segment is at least 400 km in length, ∼100 km north of the ridge–ridge–ridge triple junction between the Cocos, Pacific, and Nazca plates. Ocean crust formed at a superfast 500 km
Seafloor age (my) 0
5
10
15
20
20°N
472
100°W
159
erton
Clipp
10°N
495 80°W
FZ
845 844
Site 1256 852/1243 851
79
Hess Deep
83/503
Cocos plate
82
81/572 80
504 Pacific plate
Nazca plate
FIGURE 4.2.1.11 Age map of the Cocos plate and corresponding regions of the Pacific and Nazca plates (Wilson et al., 2003). Isochrons at 5 my intervals have been converted from magnetic anomaly identifications according to the timescale of Cande and Kent (1995). Selected DSDP and ODP sites that reached basement are indicated by circles. The wide spacing of the 10–20 my isochrons to the south reflects the extremely fast (200–220 mm per year) full spreading rate.
Formation and Evolution of Oceanic Lithosphere Chapter | 4.2.1 479
spreading rate was deliberately targeted because there is strong evidence from mid-ocean ridge seismic experiments that gabbros occur at shallower depths in intact ocean crust with higher spreading rates (Carbotte, Mutter, Mutter, & Ponce-Correa, 1998; Purdy, Kong, Christeson, & Solomon, 1992; Wilson et al., 2006). Consequently, the often difficult-to-drill upper ocean crust should be relatively thin. Site 1256 has a seismic structure reminiscent of typical Pacific offaxis seafloor (Figure 4.2.1.12(B)). Upper Layer 2 velocities are 4.5–5 km/s, and the Layer 2–3 transition is between ∼1200 and 1500 m subbasement. The total crustal thickness at Site 1256 is estimated at approximately 5–5.5 km. Site 1256 sits atop a region of smooth seafloor and basement topography (<10 m relief; Figure 4.2.1.12). Unpublished processing (A.J. Harding, personal communication; Teagle et al., 2006) reveals discernible variation in the average seismic velocity (approximately 4.54–4.88 km/s) of the uppermost (∼100 m) basement. Two principal features are apparent (Figure 4.2.1.12(A)): a 5–10-km-wide zone of relatively high upper basement velocities (>4.82 km/s) that can be traced ∼20 km to the edge of data coverage southeast of Site 1256, and a bull’s eye of relatively low velocity (4.66–4.54 km/s) centered around the crossing point of two seismic Lines (21 and 25; Figure 4.2.1.12). The only serious drawback to the Site 1256 area as a crustal reference section for fast spreading rates is its low paleolatitude, making the determination of magnetic polarity from azimuthally unoriented core samples very challenging. In addition, the nearly north–south ridge orientation at the time of accretion makes the magnetic inclination relatively insensitive to tilting.
4.2.1.3.2 Summary of Upper Crustal Lithology at Site 1256 Hole 1256D was the first scientific borehole prepared for deep drilling in ocean crust by installing a large reentry cone secured with almost 270 m of 16-inch casing through the 250-m-thick sedimentary overburden and cemented into the uppermost basement. It was then deepened through an ∼810-m-thick sequence of lavas and a thin (∼346 m) sheeted dike complex, the lower 60 m of which is strongly contact metamorphosed to granoblastic textures. The first gabbroic rocks were encountered at 1407 mbsf (IODP Expedition 312), where the hole entered a complex dike–gabbro transition zone that includes two 20–50-m-thick gabbro lenses intruded into dikes with granoblastic textures (Figures 4.2.1.13 and 4.2.1.14). At the end of IODP Expedition 312, Hole 1256D had a total depth of 1507 mbsf and was open to its full depth. Five and a half years later, IODP Expedition 335 aimed to deepen Hole 1256D several hundred meters into the cumulate gabbroic rocks of intact lower oceanic crust to address fundamental scientific questions about lower crustal accretion and cooling mechanisms. Unfortunately Hole 1256D was deepened only minimally (to 1521 mbsf), as a number of significant engineering challenges were encountered and overcome during the cruise, each unique but not unexpected in a deep, uncased marine borehole into igneous rocks (see Operations section of the Expedition Summary in Teagle et al., 2012).
(B)
FIGURE 4.2.1.12 Characteristics of the Site 1256 area (modified from Teagle et al., 2012). (A) Bathymetry; abyssal hill relief of as much as 100 m is apparent in the southwest part of the area; relief to the northeast is lower and less organized. Numbers 21–28 identify multichannel seismic (MCS) lines (white) from the site survey. Red (light gray in print versions) and blue (dark gray in print versions) contours are selected top-of-basement seismic velocity contours; the red contour encloses velocity >4.82 km/s, which is interpret as a plausible proxy for the presence of thick ponded lava flows, as sampled in Hole 1256D (Teagle et al., 2012; Wilson et al., 2006); the blue contour encloses velocities <4.60 km/s, possibly reflecting a greater portion of pillow lavas than elsewhere in the region. (B) Site survey MCS profile (Hallenborg et al., 2003) from line 22 that crosses at Site 1256, with penetration as of the end of Expedition 335 scaled approximately to traveltime. Horizontal reflectors in the upper basement to traveltimes of 5.5 s appear to result from contrasts between lava flow sequences, corresponding to depths of at least 800 m in basement.
480 Earth and Life Processes Discovered from Subseafloor Environments
(A)
Formation and Evolution of Oceanic Lithosphere Chapter | 4.2.1 481
(A)
Lith. Sedimentary overburden
0
Recovery
Massive basalt Sheet flow Pillow basalt Sheeted dike Granoblastic dike Gabbro
40
250
Resistivity ( m)
1
10 102 103 104 105
(C)
0
Porosity (%) 5
10
15 4
(D)
Vp (km/s) 5
6
7
(E)
Inflated flows
Sheet and massive flows
Leg 206
70
Refraction velocity measurement Inversion of traveltimes from Leg 206 VSP Inversion of traveltimes from Expedition 312 VSP
Exp 312
1250
Sheeted dike complex
Exp 309
Depth (mbsf)
750
1000
60
(B)
Lava pond
500
Mg #
50
Intrusive contacts Magmatic contacts Breccia
1500 Junk Baskets
FIGURE 4.2.1.13 Downhole plots of data for ODP Hole 1256D (modified from Wilson et al., 2006). (A) Summary lithostratigraphic column of the basement drilled at Site 1256 showing recovery, and major lithologies. (B) Mg# obtained from shipboard and published whole rock measurements (Neo, Yamazaki, Miyashita, & the Expedition 309/312 Scientists, 2009; Teagle et al., 2006, 2012; Wilson et al., 2003; Yamasaki et al., 2009). (C) Electrical resistivity logging (deep) data (Teagle et al., 2012). (D) Porosity (Gilbert & Salisbury, 2011). Small black symbols are shipboard measurements on discrete samples, red (dark gray in print versions) symbols are postcruise laboratory measurements, and the light blue (pale gray in print versions) curve is calculated from the DLL (Dual Laterolog tool) deep electrical resistivity log. (E) Compressional seismic velocity (modified from Gilbert & Salisbury, 2011). Small black symbols are shipboard measurements on discrete samples corrected for lithostatic pressure, red (dark gray in print versions) symbols are postcruise laboratory measurements under lithostatic pressure, the blue (light gray in print versions) curve is the sonic log data, as reprocessed by Guérin, Goldberg, and Iturrino (2008), the dashed red (dark gray in print versions) line is derived from seismic refraction data (Wilson et al., 2003), the bold red (dark gray in print versions) and orange (light gray in print versions) lines show the inversion of traveltimes from Leg 206 and Expedition 312 vertical seismic profiles (VSPs), respectively (Swift et al., 2008).
However, IODP Expedition 335, returned a unique collection of samples from the granoblastic dikes (Figures 4.2.1.13 and 4.2.1.14), obtained from the many fishing, reaming, and cleaning runs with junk baskets. Some of these samples were much larger than normal core pieces, up to ∼20 cm in equivalent diameter.
(A)
Olivine gabbro
Gabbro
Gabbronorite
100%
Oxide gabbro
1400
0%
Diorite/Tonalite
Recovery
Granoblastic dike
482 Earth and Life Processes Discovered from Subseafloor Environments
(B)
Lith. 205R-1, 10-14 cm; 1382 mbsf
(C)
1410 1420
213R-1, 52-60 cm; 1407 mbsf 214R-1, 48-58 cm; 1411 mbsf Gabbro 1
1430
(D) (E)
214R-2, 61-70 cm; 1413 mbsf
1450 Exp 312
1460
Olivine Gabbro
Gabbro
223R-2, 64-75 cm; 1451 mbsf Dike Screen 1
Depth (mbsf )
1440
1470 1480
(F)
(G)
Dike enclave
2 1
Gabbro
Gabbro 2
232R-2, 32-46 cm; 1494 mbsf 236R-1, 38-39 cm; 1512 mbsf
(H)
335-Run12-RCJB-RockB Junk Basket
Diorite/Tonalite Late dike
1521.6mbsf
Diorite vein
Basalt enclaves Dikes
(J)
(I)
ont act
1520
Diorite patch
ec
Exp 335
1510
Dike Screen 2
1500
Dik
1490
Oxide gabbro Gabbronorite Gabbro Olivine gabbro
335-Run20-RCJB-RockC Junk Basket
335-Run12-RCJB-RockS Junk Basket
FIGURE 4.2.1.14 Plutonic section from the lower portion of Hole 1256D with representative photomicrographs of key samples (modified from Wilson et al., 2006; Teagle et al., 2012). The distribution of rock types is expanded proportionately in zones of incomplete recovery. (A) Photomicrograph of a dike completely recrystallized to a granoblastic association of equant secondary plagioclase, clinopyroxene, magnetite, and ilmenite (plane-polarized light). (B) Uppermost dike/gabbro boundary; medium-grained oxide gabbro is intruded into a granoblastically recrystallized dike along an irregular, moderately dipping contact. (C) Quartz-rich oxide diorite strongly altered to actinolitic hornblende, secondary plagioclase, epidote, and chlorite. (D) Disseminated oxide gabbro with patchy texture and centimeter-scale dark ophitically intergrown clinopyroxene and plagioclase patches separated by irregular, more highly altered leucocratic zones. (E) Sharp modal contact between a medium-grained olivine-rich gabbro (top in the core) and a gabbro. (F) Clast of partially resorbed granoblastic dike within gabbro. (G) Photomicrograph of a granoblastic basalt (plane-polarized light) showing a relict plagioclase microphenocryst with pyroxenes and oxide inclusions in its core (1) and a metamorphic, orthopyroxene vein (2). (H) Dioritic vein and
Formation and Evolution of Oceanic Lithosphere Chapter | 4.2.1 483
The volcanic sequences at Site 1256 consist of massive flows and thinner sheet flows with subordinate pillow basalt, hyaloclastite, and breccia ( Figure 4.2.1.13). The uppermost crust at Site 1256 comprises an ∼100-m-thick sequence of lava dominated by a single massive lava flow >75-m thick, requiring at least this much seafloor relief to pond the lava (Tartarotti & Crispini, 2006; Wilson et al., 2003). On modern fast-spreading ridges, such topography does not commonly develop until 5–10 km from the axis (Macdonald, Fox, Alexander, Pockalny, & Gente, 1996). Although this lava flow cooled off-axis it may have originated at the ridge axis before flowing on to the ridge flanks, as is observed for very large lava flows on modern ocean floor (e.g., Macdonald, Haymon, & Shor, 1989). This flow is relatively unfractured at the scale of the drill hole (Crispini, Tartarotti, & Umino, 2006), with shipboard physical property measurements on discrete samples indicating VP > 5.5 km/s (Wilson et al., 2003). Hence, the area of relatively high uppermost basement seismic velocities likely represents the regional extent of this massive flow, although this hypothesis has not been tested by drilling. Assuming an average thickness of 40 m, the regional survey results would conservatively suggest an eruption volume in excess of 3 km3, plausibly >10 km3. Lava flows this voluminous are not described on-axis at fast-spreading ridges. An off-axis origin is suggested by the presence of a strongly welded and plastically deformed agglutinate in the base of the thick lava in Hole 1256C, that indicates proximity to the source vent for the lava field (Crispini et al., 2006). An off-axis volcanic flow field of comparable size is described on the EPR at 14°S (Geshi et al., 2007). This is extremely large when compared to the size of mid-ocean ridge axial lowvelocity zones that are thought to be high-level melt lenses, which typically have volumes approximately 0.05–0.15 km3 per kilometer of ridge axis, and generally appear to be only partially molten (Singh, Kent, Collier, Harding, & Orcutt, 1998). This poses questions regarding the melt sources of these very large submarine flow fields that require either the draining of a relatively large magma reservoir so far not imaged at fast-spreading mid-ocean ridges (Teagle et al., 2012), or a magma system that is being recharged during the course of slow (several decades long), ongoing eruptions, such as described in Iceland (e.g., Eason & Sinton, 2009). The lava sequence below the lava pond includes sheet and massive flows and minor pillow flows. Subvertical, elongate flow-top fractures filled with quenched glass and hyaloclastite in these lavas indicate flow lobe inflation requiring cooling on a subhorizontal surface off-axis (Umino, Lipman, & Obata, 2000). It is estimated that a total thickness of 284 m of lava flowed and cooled off-axis. Sheet
patch within granoblastic basalt. (I) Sharp, curviplanar dike/dike contact between a light-colored, fine-grained basalt (right) and a darker, coarse-grained basalt (left). (J) Medium-grained, orthopyroxene-bearing olivine gabbro.
484 Earth and Life Processes Discovered from Subseafloor Environments
flows and massive lava flows erupted at the ridge axis make up the remaining extrusive section to 1004 mbsf. This contrasts slightly with the volcanic stratigraphy for Hole 1256D developed from analysis of wireline geophysical imaging (Tominaga & Umino, 2010; Tominaga, Teagle, Alt, & Umino, 2009) that suggests <50% of the lava drilled crystallized within 1000 m of the axis but that the majority of the lava pile had formed within 3000 m of the ridge crest. From 1004 to 1061 mbsf is the Transition Zone between the extrusive rocks and the sheeted dike complex, marked by the occurrence of subvertical intrusive contacts, a silica-sulfide mineralized hyaloclastic breccia zone, and a step change in hydrothermal alteration from low temperature to greenschist facies assemblages. Below this transition, numerous subvertical intrusive contacts were observed, indicating the start of a relatively thin, ∼350-m-thick sheeted dike complex dominated by massive basalt. Some basalts have doleritic textures, and many are crosscut by subvertical dikes with strongly brecciated chilled margins that are mineralized by hydrothermal fluids. There is little evidence from core or geophysical wireline logs for major tilting of the dikes, consistent with subhorizontal seismic reflectors in the lower extrusive rocks that are continuous for several kilometers across the site (Hallenborg, Harding, Kent, & Wilson, 2003). Measurements of dike orientations using FMS images are in close agreement with direct measurements on recovered cores and indicate that the dikes at Site 1256 are tilted ∼15° away from the paleospreading axis around a ridgeparallel, subhorizontal axis (Tominaga et al., 2009; Violay, Pezard, Ildefonse, Célérier, & Deleau, 2012). In the bottommost ∼60 m of the sheeted dikes (1348–1407 mbsf), basalts are partially to completely recrystallized to distinctive granoblastic textures characterized by granular secondary clinopyroxene and lesser orthopyroxene, resulting from contact metamorphism by underlying gabbroic intrusions (Figures 4.2.1.13 and 4.2.1.14). Textural changes in oxide minerals indicate that the zone of metamorphic recrystallization may extend for >90 m above the upper contact with the first gabbro interval (Coggon, Alt, & Teagle, 2008). There is unequivocal evidence for hydrothermal alteration before and after the formation of the granoblastic textures and potentially small, irregular patches of partial melting (Alt et al., 2010; France, Ildefonse, & Koepke, 2009; Koepke et al., 2008; Teagle et al., 2006). Simple thermal calculations suggest that the two gabbro bodies so far intersected in Hole 1256D have insufficient thermal mass to be responsible for a 60–90-m-thick high-temperature (600–900 °C) contact-metamorphic halo if the intrusions are simple subhorizontal bodies (Coggon et al., 2008; Koepke et al., 2008). The amount of heat needed to produce this metamorphism requires either a much larger intrusion nearby, or significant topography on the dike/gabbro boundary (Teagle et al., 2006). Gabbro and trondhjemite intrusions into sheeted dikes at 1407 mbsf mark the top of the plutonic section (Figures 4.2.1.13 and 4.2.1.14) and the start of a complex dike–gabbro transition zone. This first occurrence of plutonic rocks from intact ocean crust occurred within the depth range predicted from
Formation and Evolution of Oceanic Lithosphere Chapter | 4.2.1 485
the relationship between spreading rate and depth to axial low-velocity zones (Teagle et al., 2006). Two major bodies of gabbro were penetrated beneath this contact, with the 52-m-thick upper gabbro (Gabbro 1) separated from the 24-m-thick lower gabbro (Gabbro 2) by a 24-m screen of granoblastic basalts (Dike Screen 1). The textures and rock types observed in Hole 1256D gabbro intervals are similar to those of varitextured gabbros thought to represent a frozen melt lens between the sheeted dike complex and the underlying cumulate gabbros in ophiolites (e.g., MacLeod and Yaouancq, 2000; France et al., 2009). Gabbro 1 is mineralogically and texturally heterogeneous, comprising olivine gabbros, gabbros, oxide gabbros, quartz-rich oxide diorites, and small trondhjemite dikelets. Oxide abundance decreases irregularly downhole, and olivine is present in significant amounts in the lower part of Gabbro 1. The intervening Dike Screen 1 is interpreted as an interval of sheeted dikes captured between the two intrusions of gabbros. It consists of fine-grained meta-basalt similar to the granoblastic dikes overlying Gabbro 1, and is cut by a number of small quartz gabbro and tonalite dikelets of variable thickness (1–10 cm), grain size, and composition. Gabbro 2 comprises gabbro, oxide gabbro, and subordinate orthopyroxene-bearing gabbro and trondhjemite that are altered similarly to Gabbro 1, and has clear intrusive contacts with the overlying granoblastic Dike Screen 1. Partially resorbed, stoped dike clasts are entrained within both the upper and lower margins of Gabbro 2 (Figure 4.2.1.14). Gabbro 2 is characterized by the absence of fresh olivine, high but variable orthopyroxene contents (5–25%), and considerable centimeter- to decimeter-scale heterogeneity. Oxide abundance generally diminishes downhole. The predominant rock type is orthopyroxene-bearing gabbro with gabbronorite in the marginal units. Dike Screen 2, at the bottom of the hole, was extensively sampled with junk baskets during IODP Expedition 335 (Teagle et al., 2012). It is dominantly made of fine-grained granoblastic aphyric basalts, most of which are strongly recrystallized (Figure 4.2.1.14). Some samples preserve dike-to-dike contacts, and many preserve cores of former plagioclase (micro)phenocrysts (Figure 4.2.1.14), indicating that the granoblastic rocks are contact-metamorphosed sheeted dike diabase. Approximately half of the granoblastic basalt samples contain small, irregular patches (<5 × 3 cm), veins (approximately 1–2 mm wide), and dikelets (<1.5 cm wide) of evolved plutonic rocks (oxide gabbro, diorite, and tonalite) (Figures 4.2.1.14 and 4.2.1.15). The veins are observed to be offshoots of the igneous patches, and diffuse patches are issued from the dikelets. Hence these features form part of the same network, marking a single generation of intrusion of melts into the granoblastic basalts. Their magmatic textures (subhedral to euhedral shapes of several phases, along with poikilitic textures) contrast strongly with the granoblastic textures of the host rocks, demonstrating that intrusion occurred after the high-temperature recrystallization of the host rock. The common occurrence of primary magmatic amphibole in the veins and patches suggests high water activities during their formation, hence possibly the partial
486 Earth and Life Processes Discovered from Subseafloor Environments
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FIGURE 4.2.1.15 Schematic “outcrop” of the dike–gabbro transition zone, showing the micro-textures and contact relationships observed in the cores and junk basket samples recovered during Expedition 335 (Teagle et al., 2012). (A) Diorite dikelet, comprised predominantly of primary igneous amphibole and plagioclase, crosscutting granoblastic basalt (Sample 335-1256D-Run19-RCJB-Rock C). (B) Fine-grained dike chilled against coarse-grained dike (Sample 335-1256D-Run12-RCJB-Rock S). The entire sample is recrystallized to granoblastic assemblages of plagioclase, orthopyroxene, clinopyroxene, magnetite, and ilmenite. Later fine postcontact-metamorphic hydrothermal amphibole veins (not shown) cut across the contact and both dikes. (C) Chilled and brecciated dike margin recrystallized to granoblastic assemblages (Sample 335-1256D-Run14-EXJB-Foliated). Angular clasts consist of granoblastic plagioclase with minor orthopyroxene, clinopyroxene, ilmenite, and magnetite and are recrystallized from chilled dike margin breccia protolith. Clast matrix is orthopyroxene-rich with minor clinopyroxene, plagioclase, and oxides; it is recrystallized from hydrothermal minerals (amphibole and chlorite) that veined and cemented the breccia protolith. (D) Subophitic texture in gabbro (Sample 335-1256D-Run11EXJB). (E) Diorite vein crosscutting a conjugate set of metamorphic orthopyroxene veins (Sample 335-1256D-Run19-RCJB-Rock B). (F) Postcontact-metamorphic hydrothermal hornblende vein cutting granoblastic basalt (Sample 335-1256D-238R-1, 2–4 cm). (G) Granoblastic basalt with a diorite patch (Sample 335-1256D-Run11-EXJB-Rock). (H) Granoblastic orthopyroxene vein, recrystallized from precursor hydrothermal vein, in granoblastic dike (Sample 335-1256D-238R-1, 2–4 cm). Orthopyroxene vein is cut by small postcontact-metamorphic hydrothermal amphibole vein. (I) Postcontact-metamorphic hydrothermal amphibole vein cutting granoblastic basalt, with 1-mm-wide amphibole-rich alteration halo where pyroxenes are replaced by amphibole (Sample 335-1256D-Run12-RCJB-Rock B).
melting of previously hydrothermally altered rocks. Moreover, the presence of quartz, as well as accessory apatite and zircon, implies that patches, veins, and dikelets crystallized from highly evolved melts. Only a small number of gabbroic rocks (∼3 wt%) were recovered during IODP Expedition 335
Formation and Evolution of Oceanic Lithosphere Chapter | 4.2.1 487
(Figure 4.2.1.14) and these rocks range in composition from disseminated oxide gabbro to orthopyroxene-bearing olivine gabbro. The IODP Expedition 335 gabbroic rocks are texturally less variable than those recovered during IODP Expedition 312 from the gabbro 1 and 2 intervals. Although the contact relationships with the granoblastic basalts were not recovered, it is likely that at least some of the gabbroic rocks occur intercalated with the granoblastic basalt, perhaps forming small intrusions. The unique sampling with fishing tools achieved during IODP Expedition 335 allowed the documentation of a variety of intrusive structural/textural relationships, and overprinting/crosscutting hydrothermal alteration and metamorphic paragenetic sequences that had not previously been observed in situ, owing to the one-dimensional nature of drill cores and the very low rates of recovery of the granoblastic material in the sheeted dike/gabbro transition zone. Overall, the bottom of Hole 1256D documents a section of metamorphosed, granoblastic sheeted dikes that underwent small-scale intrusion by both gabbroic and evolved plutonic rocks (Figure 4.2.1.15).
4.2.1.3.3 A Thin upper Crust at Superfast Spreading Rate Relative to other well-studied upper ocean crust sections (e.g., Karson, 2002), Site 1256 is unique as it shows a thick lava sequence and a thin dike sequence (Umino et al., 2008). Steady-state thermal models require that the conductive lid separating magma from rapidly circulating seawater gets thinner as spreading rate increases, indicating that the thin dike sequence is a direct consequence of the high spreading rate. A thick flow sequence, with many massive individual flows and few pillow lavas, is a consequence of the elevated magma budget (Umino et al., 2008) and/or of the short vertical transport distance from the magma chamber, and is similar to observations from the midsegments of the fast-spreading southern EPR (White, Macdonald, & Haymon, 2000). The high lava/dike thickness ratio appears to be characteristic of superfast spreading, which is in direct contrast to spreading models developed from observations of tectonically disrupted fast-spread crust exposed in Hess Deep (Karson, 2002) that suggest regions of high magma supply should have thin lavas and thick dikes. Similarly, there is little evidence for tilting (at most ∼15°) in Hole 1256D and no evidence for significant faulting. In contrast, the upper crust exposed at Hess Deep (fast spreading) and in the Blanco Fracture Zone (intermediate spreading) shows significant faulting and rotation within the dike complex (Karson, 2002), indicating that observations from those tectonic windows may be site specific and not widely applicable to intact ocean crust. The massive lava flow at the top of the Site 1256 basement indicates that faults with throws of approximately 50–100 m must exist in superfast spreading rate crust relatively near to the ridge axis (<10 km) in order to provide the relief necessary to pond the lava (as described at the southern EPR; e.g., Auzende et al., 1996; Macdonald, 1998; Sinton et al., 2002).
488 Earth and Life Processes Discovered from Subseafloor Environments
4.2.1.3.4 Where Is the Layer 2/3 Transition at Site 1256? Marine seismologists traditionally subdivided the ocean crust into seismic layers: Layer 1 comprises low-velocity sediments (VP < 3 km/s); Layer 2 has low velocity and a high velocity gradient, with VP typically ranging from ∼3.5 to ∼6.7 km/s; and Layer 3 has high velocity and a low velocity gradient (VP ranges from 6.7 to 7.1 km/s); the step-up to VP > 8 km/s marking the Mohorovičić Discontinuity and is interpreted to be the boundary with ultramafic rocks of the uppermost mantle. In casual language “Layer 3” is often mistakenly used as a synonym for gabbro, although previous drilling in Hole 504B has penetrated Layer 3 but not yet reached gabbroic rocks (Alt et al., 1996; Detrick et al., 1994). From regional seismic refraction data, the transition from seismic Layer 2 to Layer 3 at Site 1256 occurs between 1200 and 1500 m into basement (Wilson et al., 2003). Shipboard determinations of seismic velocities of discrete samples, once corrected for lithostatic pressure, are in close agreement with postcruise laboratory measurements of discrete samples under lithostatic pressure, and with in situ measurements by wireline tools (Figure 4.2.1.13; Gilbert & Salisbury, 2011). Contrary to expectation, porosity increases and P-wave velocities decrease stepwise downward from the lowermost dikes into the uppermost gabbro in Hole 1256D, as the result of the contact metamorphism of the granoblastic dikes and the strong hydrothermal alteration of the uppermost gabbros (Figure 4.2.1.13). Porosity and velocity then increase downhole in the gabbro. Wireline velocity measurements end at the top of gabbro, because additional velocity data closer to the bottom of the hole could not be acquired at the end of IODP Expedition 335 (Teagle et al., 2012). The examination of shipboard and postcruise discrete sample measurements, wireline logging, and VSP data suggests that the base of Hole 1256D has just entered, or is close to the transition from Layer 2 to Layer 3 (Gilbert & Salisbury, 2011; Swift, Reichow, Tikku, Tominaga, & Gilbert, 2008). Encountering gabbro at a depth within Layer 2 reinforces previous inferences that factors such as porosity and hydrothermal alteration (Alt et al., 1996; Carlson, 2010; Detrick et al., 1994) are more important than rock type or grain size in controlling the location of the Layer 2–3 transition. Drilling deeper in Hole 1256D is required to characterize the Layer 2–3 transition at Site 1256.
4.2.1.3.5 Geochemical Characteristics of the Upper Crust at Site 1256 Flows and dikes from Hole 1256D are N-MORB but show a wide range of magmatic fractionation (Figures 4.2.1.13 and 4.2.1.16). The lavas and dikes are on average relatively fractionated compared to global MORB data compilations as anticipated for fast-spreading EPR MORB (e.g., Rubin & Sinton, 2007). The lava pond, cored in both Holes 1256C and D, is more fractionated (Mg# 50 ± 2) than the rocks from the inflated flows, sheet and massive flows, and the sheeted
Formation and Evolution of Oceanic Lithosphere Chapter | 4.2.1 489 6KHHWIORZ
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dike complex that all show a similar range of compositions (Mg# 56 ± 6). The more differentiated nature of the lava pond lavas is also displayed by relatively high incompatible element (e.g., Zr, TiO2, Y, and V) and low compatible element (Cr and Ni) concentrations (Teagle et al., 2006; Wilson et al., 2003). An interval within the lava pond cored in Hole 1256C (294–306 mbsf) shows anomalously high K2O concentrations (∼0.74 wt%) that are difficult to explain by magmatic fractionation and are not a result of hydrothermal alteration by seawater. There are variations in the basalt chemistry downhole, with a step in igneous trends at 750 mbsf possibly indicating a cycle of fractionation, replenishment, and, perhaps, assimilation (Figure 4.2.1.13). Downhole geochemical compositions within the dike section are variable and do not define trends; primary and evolved compositions are closely juxtaposed, as would be expected for vertically intruded magmas (Figure 4.2.1.13). The range of major element compositions in the dikes is similar to that of the overlying rocks, and the average composition of the dikes is indistinguishable from the average composition of the lavas (Figures 4.2.1.13 and 4.2.1.16). The compositional ranges of fresh lava and dike samples correspond to typical values for N-MORB for most major elements and many trace elements (Gale, Dalton, Langmuir, Su, & Schilling, 2013; Rubin & Sinton, 2007) and are
490 Earth and Life Processes Discovered from Subseafloor Environments
similar to those observed for the northern EPR (Figure 4.2.1.16). When trace elements are compared to EPR MORB, they are within one standard deviation of average, albeit on the relatively trace element-depleted side of MORB. For example, compared with first-order mid-ocean-ridge segments along the EPR, basalts from Site 1256 have low Zr/TiO2 and Zr/Y. This decrease in trace element ratios would require ∼30% more melting at the superfast ridge, but this appears unlikely because of normal crustal thickness (∼5.5 km) at Site 1256. More likely, the relatively depleted trace element signature at ODP Site 1256 results from the upwelling of previously depleted Galápagos plume mantle at the paleo-Site 1256 ridge axis (Geldmacher et al., 2013; Park, MacLennan, Teagle, & Hauff, 2008). The gabbro compositions, as documented by shipboard geochemical analyses, span a range similar to the flows and dikes but have on average higher MgO and lower FeO, albeit still within the range of EPR basalts (Figure 4.2.1.16). The uppermost gabbros have geochemical characteristics similar to the overlying dikes with MORB chemistries (MgO approximately 7–8 wt%; Zr approximately 47–65 ppm). Deeper, in Gabbro 2, the rocks are less fractionated and there are general downhole trends of increasing MgO, CaO, and Ni and decreasing FeO, Zr, and Y. Decreasing concentrations of FeO and TiO2 downhole suggest that Gabbro 2 is fractionated similarly to Gabbro 1. The intrusive nature of both gabbro bodies and the chemical variations within them suggest that they intruded into the base of the sheeted dikes and underwent minor internal fractionation in situ, resulting in the observed general geochemical stratification. Partially resorbed xenoliths of granoblastic dikes within Gabbro 2 (Figure 4.2.1.14) indicate that stoping of the intruded dikes may have contaminated the gabbro compositions (e.g., France et al., 2009). There are linear trends in the gabbro units between MgO and TiO2, FeO, CaO, Na2O, and Zr, most likely resulting from fractional crystallization of a gabbroic magma, with the fractionating assemblage consisting of clinopyroxene, plagioclase, and subordinate olivine, as expected for relatively evolved basaltic magmas. Olivines within the gabbros have Forsterite composition Fo < 81 (Yamazaki, Neo, & Miyashita, 2009) indicating fractionation. Similarly, even the most primitive dikes and lavas at Site 1256 (Mg# ∼66) would be in equilibrium with olivine more evolved than mantle olivine, and hence must have undergone fractionation. Simple mass balance calculations indicate that the average basalt in Hole 1256D has lost >30% of its original liquid mass as solid gabbro, implying the presence of at least 300 m of cumulate gabbro in the crust below the present base of the hole (Teagle et al., 2006). Therefore, the residue removed from primary magma to produce the observed gabbro and basalt compositions must occur below the present base of Hole 1256D. The lower part of Hole 1256D, below the dike–Gabbro 1 boundary at ∼1406 mbsf, is characterized by large chemical variations with, for example, Mg# ranging from 42 to 72 (Figure 4.2.1.13), and Zr from 23 to 117 ppm (Teagle et al., 2012). This variability mainly reflects changes in rock types from the
Formation and Evolution of Oceanic Lithosphere Chapter | 4.2.1 491
low-Mg#, trace element-rich sheeted and granoblastic dikes and dike screens, to the higher Mg# and trace element-depleted gabbroic rocks of Gabbro 1 and Gabbro 2. There are significant depth-dependent trace element variations in the granoblastic basalts, with a general downhole trend of decreasing incompatible element contents (i.e., Zr and, to a lesser extent, Y). Details of these variations and their possible significance are presented in Teagle et al. (2012).
4.2.1.4 SHALLOW DRILLING IN FAST-SPREAD LOWER CRUST AT HESS DEEP The primary goal of the “superfast campaign” at Site 1256 was to reach gabbroic rocks of the lower crust. In spite of the operational difficulties encountered during IODP Expedition 335, Hole 1256D has been thoroughly cleaned from debris and cuttings, and is still open. Additional expeditions to Hole 1256D will be required to deepen the hole into the cumulate rocks of the lower crust. An alternative to deep drilling into intact, fast-spread oceanic crust (as at Site 1256) is to access the lower crust in the rare places where it crops out on the seafloor, such as in tectonic windows. This was the goal of the recent IODP Expedition 345, which targeted lower crust gabbros at Hess Deep in the equatorial eastern Pacific (Figure 4.2.1.1; Table 4.2.1.1; Gillis, Snow, Klaus, & the Expedition 345 Scientists, 2014a; Gillis et al., 2014b). The scientific objectives were to test endmember models of magmatic accretion and estimate the intensity of hydrothermal cooling at depth. The Hess Deep Rift was selected to exploit tectonic exposures of young EPR plutonic crust, building upon results from Ocean Drilling Program Leg 147 (Gillis et al., 1993) and submersible surveys (e.g., Francheteau et al., 1990; Hékinian, Bideau, Francheteau, Lonsdale, & Blum, 1993; Lissenberg, MacLeod, Howard, & Godard, 2013). Low to moderate recovery (15–30%) was achieved in three holes (35–110 mbsf). Olivine gabbro and troctolite represent 84% of the plutonic rock types recovered at Site U1415. Additional recovered lithologies are minor gabbro, clinopyroxene oikocryst-bearing troctolite, clinopyroxene oikocryst-bearing gabbro, and gabbronorite. All of these lithologies have Mg# between 78 and 90. Spectacular modal and/or grain-size layering, reminiscent of that classically described in ophiolites (e.g., Boudier, Nicolas, & Ildefonse, 1996; Kelemen, Koga, & Shimizu, 1997; Pallister & Hopson, 1981), was observed in more than half of the recovered material (Figure 4.2.1.17). Magmatic foliation is commonly moderate to strong, particularly in layered lithologies. Together with the compositions of gabbros from the upper part of the igneous crust (known from ODP Hole 894 samples; Gillis et al., 1993), and of the upper crust (known from numerous Alvin and Nautile dive samples; Coogan, Gillis, MacLeod, Thompson, & Hekinian, 2002, and references therein), Gillis et al. (2014b) calculated an estimate of the bulk composition of the Hess Deep crust, the first estimate for fast-spread crust, with an Mg# of 74. This average bulk composition can then be used as that of a parental melt for a MORB fractional crystallization model, which is consistent with the measured sample compositions (Gillis et al., 2014b). However, an unexpected finding, which
492 Earth and Life Processes Discovered from Subseafloor Environments
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does not fit a simple fractional crystallization trend, is the common abundance (up to 5%) of orthopyroxene as an early cumulus phase in the rocks recovered during Expedition 345. This observation calls for a more complex petrogenetic history of the Hess Deep lower crust, which possibly involves reaction of ascending MORB melts with the upper mantle (hence delivery of melts to the lower crust that are not undersaturated in orthopyoxene; Coogan et al., 2002), and/or the contamination of parental melts by water (Boudier, Godard, & Armbruster, 2000; Nonnotte, Ceuleneer, & Benoit, 2005). The common occurrence of delicate skeletal olivine grains suggests that at this location, the lower crust was not significantly deformed after crystallization. IODP Expedition 345 provides a new, unique, reference section for fast-spreading lower crust (Gillis et al., 2014b), which complements the previous seafloor sampling in this area (Hékinian et al., 1993; Coogan et al., 2002; Lissenberg et al., 2013). Paleomagnetic demagnetization experiments conducted during IODP Expedition 345 on discrete samples of gabbros and troctolites yielded stable magnetization directions with a variety of remanence structures. In all samples, components that unblock close to the magnetite Curie temperature are considered to represent primary thermoremanent magnetizations acquired during crustal accretion. Downhole variations in the inclination of this stable, highunblocking temperature magnetic component indicate that at least two blocks have been sampled in each of the two deepest holes (U1415J and U1415P; Gillis, Snow,
Formation and Evolution of Oceanic Lithosphere Chapter | 4.2.1 493
Klaus, & the Expedition 345 Scientists, 2014a). The data provide evidence of relative displacements of individual, internally coherent blocks (tens of meters in size), consistent with emplacement by very large-scale mass wasting on the southern slope of the Hess Deep intrarift ridge (Ferrini et al., 2013).
4.2.1.5 CONCLUSION During the last decade, IODP has made substantial progress in sampling of the ocean crust, by doubling the amount of deep (>1 km below seafloor) holes. Two deep holes from sections formed at extremely different spreading rates have shed light on very important unknown and/or poorly understood aspects of the architecture and accretion of modern ocean crust. In slow-spread crust, Hole U1309D was the second deep hole in an OCC following historic deep drilling at ODP Hole 735B (that remains the deepest hole in slow-spread crust, and returned the highest recovery in hard rock drilling). The spectacular igneous section recovered in the core of the Atlantis Massif is key to demonstrating the important role of magmatism in the development of OCCs. The interplay between magmatism and tectonics is still to be fully documented, described, and quantified, but Hole U1309D is a significant step forward. Olivine-rich troctolites, which represent about 5% of the material recovered in the hole, and which are by far the least altered olivine-rich material (locally <1% serpentinized) ever recovered by scientific ocean drilling, represent an exceptional record of the complex igneous accretion history, involving both fractional crystallization and melt–rock reactions, and possibly also assimilation of mantle peridotites. Another first, resulting from the combination of high recovery and high-quality borehole imaging (FMS) data, was the core-log integration that allowed reorientation of individual core pieces and associated paleomagnetic data and the provision of unequivocal constraints on the tectonic rotation associated with the development of the OCC. Hole 1256D, in Pacific superfast-spread crust, has for the first time reached the base of the sheeted dikes and penetrated the first gabbroic rocks ever sampled in intact, in situ layered fast-spread crust. The upper crustal section recovered at Site 1256 complements Hole 504B in intermediate-spread crust, which has been for a long time the single available reference section for the modern oceanic upper crust. The sampling of the root zone of the sheeted dike complex, and underlying varitextured gabbros and recrystallized granoblastic basalts, documents for the first time the transition zone between the present-day upper and lower crust formed at fast-spreading ridges. From this unique sampling emerges a vision of a complex, dynamic thermal boundary layer zone. This region of the crust between the principally hydrothermal domain of the upper crust and the intrusive magmatic domain of the lower crust is one of evolving geological conditions. The resulting intimate coupling among temporally and spatially intercalated magmatic, hydrothermal, partial melting, intrusive, metamorphic, and retrograde processes is recorded in the recovered samples.
494 Earth and Life Processes Discovered from Subseafloor Environments
In addition to the two deep holes in slow- and ultrafast-spread crust, another step forward was recently achieved by sampling the first significant sections of layered gabbros of “Penrose”-style ocean crust (Gillis et al., 2014b), assumed to constitute most of the lower crust based on ophiolite studies since the early 1970s. The drawback of accessing the deeper crustal levels in a tectonic window like Hess Deep is that the recovered sections are tectonically out of place, and hence cannot be used to fully document the depth variations of compositions and structures in the way required to test crustal accretion models (e.g., Teagle et al., 2012). Yet, this unique sampling of the lower, layered crust brings fundamental new constraints on the average composition of fastspread crust, on the petrological processes involved in building the layered gabbros, and in transporting melt from these deeper levels to the upper crust. Hence, critical samples and data recovered by scientific drilling over the past 10 years provide significant additional constraints on understanding lower oceanic crustal processes. However, we are still far from having a complete set of direct observations regarding the accretion occurring beneath the dike layer at fast-spreading ridges, and we need to achieve deep drilling of intact, fast-spread ocean crust. Deep drilling into ocean crust has posed, and will continue to present major technical and programmatic challenges to scientific ocean drilling. Only four holes, DSDP Hole 504B, ODP Holes 735B and 1256D, and IODP Hole U1309D (Figures 4.2.1.1 and 4.2.1.2; Table 4.2.1.1), have been cored deeper than 1 km into oceanic basement, and these penetrations are arguably some of the greatest technical achievements of scientific ocean drilling. All were “hard won” multiexpedition experiments. From the experiences of drilling these holes, there are important lessons to be learned for the siting, planning, and implementation of future deep drilling of the oceanic basement, which are described in detail in the IODP Expedition 335 Proceedings (Teagle et al., 2012). Establishing the ideal location for drilling is only part of the challenge of successfully drilling moderately deep holes (2–3 km) to recover the samples and data necessary to address long-standing primary goals of scientific ocean drilling. Experience from Holes 504B and 1256D indicates that such experiments require multiple expeditions to achieve their target depths. A total of ∼500 m penetration per expedition is an upper limit for coring in the upper crust, with lesser advances and more frequent drilling challenges as these holes get deeper, and rocks metamorphosed at higher pressures and temperatures are encountered. Penetration and core recovery rates have been low to very low in the two sheeted dike complex sections drilled to date. Average rates of recovery and penetration in the dike section of Hole 1256D are 32% and 0.8 m/h, respectively. The average rate of recovery in the sheeted dike complex of Hole 504B was a miserly 11%. In contrast, experience to date suggests that gabbroic rocks can be cored relatively rapidly at high rates of recovery (e.g., Hole U1309D: penetration rate = 2 m/h; recovery ≥75%), so when the dike–gabbro transition zone is breached in Hole 1256D, solid progress through the plutonic section may be anticipated.
Formation and Evolution of Oceanic Lithosphere Chapter | 4.2.1 495
Scientific ocean drilling has little experience in casing long sections (hundreds of meters) of oceanic basement and a poor armory of underreaming tools for opening hard rock basement holes to the diameters required for the insertion of a casing. Casing hundreds of meters of a deep borehole in igneous basement would be a high risk, costly, and ship-time consuming operation that would produce no new scientific output until completed and drilling was resumed. However, it would greatly improve the stability and hydrodynamics of deep basement boreholes. A regular drilling-then-casing approach to investigate the lower oceanic fast-spread crust (target depth ∼3 km) will require a long-term commitment by the scientific ocean drilling community to a particular site and experiment, and possibly as many as 10 expeditions to complete. The possibility that even such a highly engineered approach could still fail to reach its target would have to be acknowledged and accepted. Such is the capriciousness of hard rock coring that scientific ocean drilling may have to consider new approaches if it is to ever successfully address some of the major science questions that remain unanswered after more than 50 years. There are unlikely to ever be “quick wins” with targets that require multi-expedition deep boreholes. Holes U1309D and Hole 1256D are open, stable, and can be deepened further. Drilling down to the bottom of Hole U1309D was achieved without major difficulty, and the hole condition is excellent. During IODP Expedition 335, Hole 1256D was stabilized, thoroughly cleared of cuttings (several hundred kilograms) and debris. Both holes are ready for deepening in the near future, and fundamental observations concerning the formation and evolution of the ocean crust await.
ACKNOWLEDGMENTS This research used samples and data provided by the Integrated Ocean Drilling Program (IODP). The manuscript benefited from thorough and helpful reviews from Gail Christeson and John Sinton. We thank the USIO teams and JOIDES Resolution crews for their invaluable assistance and outstanding work during IODP Expeditions 304, 305, 309, 312, 335, and 345.
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Formation and Evolution of Oceanic Lithosphere Chapter | 4.2.1 505 Wilson, D. S., Teagle, D. A. H., Alt, J. C., Banerjee, N. R., Umino, S., Miyashita, S., et al. (2006). Drilling to gabbro in intact ocean crust. Science, 312(5776), 1016–1020. doi:10.1126/science.1126090. Workman, R., & Hart, S. R. (2005). Major and trace element composition of the depleted MORB mantle (DMM). Earth and Planetary Science Letters, 231, 53–72. doi:10.1016/j. epsl.2004.12.005. Yamazaki, S., Neo, N., & Miyashita, S. (2009). Data report: whole-rock major and trace elements and mineral compositions of the sheeted dike–gabbro transition in ODP Hole 1256D. In D. A. H. Teagle, J. C. Alt, S. Umino, S. Miyashita, N. R. Banerjee, D. S. Wilson, et al. (Eds.), Proc. IODP, 309/312. Washington, DC: Integrated Ocean Drilling Program Management International, Inc. doi:10.2204/iodp.proc.309312.203.2009. Yeats, R. S., Hart, S. R., & DSDP Leg 34 Scientific Party. (1976). Initial reports of the deep sea drilling project (Vol. 34)Washington: U.S. Government Printing Office. doi:10.2973/dsdp. proc.34.1976.
Formation and Evolution of Oceanic Lithosphere: New Insights on Crustal Structure and Igneous Geochemistry from ODP/IODP Sites 1256, U1309, and U1415 Benoit Ildefonse,1,* Natsue Abe,2 Marguerite Godard,1 Antony Morris,3 Damon A.H. Teagle4 and Susumu Umino5 1Géosciences Montpellier, Université Montpellier, Montpellier, France; 2Institute for Research on Earth Evolution, Japan Agency for Marine-Earth Science and, Technology (JAMSTEC), Yokosuka, Japan; 3School of Geography, Earth and Environmental Sciences, Plymouth University, Plymouth, UK; 4National Oceanography Centre Southampton, University of Southampton, Southampton, UK; 5Department of Earth Sciences, Kanazawa University, Kanazawa, Japan *Corresponding author: E-mail: [email protected]
4.2.1.1 INTRODUCTION In March–April 1961, the drilling barge CUSS1 undertook the first scientific ocean drilling operation off Guadalupe Island, ∼240 km west of Baja California (Mexico). This expedition, beautifully reported in LIFE magazine by the novelist John Steinbeck and the renowned science photographer Fritz Goro, was the first (and eventually only) concrete manifestation of Project Mohole1. Project Mohole was an ambitious endeavor proposed initially in the late 1950s by the American Miscellaneous Society (AMSOC), an informal group of notable U.S. scientists, mostly geophysicists and oceanographers associated with the Office of Naval Research, including Harry Hess and Walter Munk. The principal aim was to drill through the oceanic crust, through the Mohorovičić Discontinuity, and to retrieve samples from Earth’s mantle (Teagle & Ildefonse, 2011). In his
1. http://www.nationalacademies.org/mohole.html. Developments in Marine Geology, Volume 7. http://dx.doi.org/10.1016/B978-0-444-62617-2.00017-7 Copyright © 2014 Elsevier B.V. All rights reserved.
449
450 Earth and Life Processes Discovered from Subseafloor Environments
book “A Hole in the Bottom of the Sea” (Bascom, 1961), Willard Bascom, Director of Project Mohole, notes that probably the first written suggestion for deep penetration down into the mantle was given by Frank Estabrook, an astrophysicist from the Basic Research branch of the U.S. Army in Pasadena (Estabrook, 1956). In addition to accomplishing a major technological breakthrough (including the invention of dynamic positioning, the drilling guide horn, and deepwater drill hole reentry), the Project Mohole drilling expedition in 1961 cored into oceanic basement for the first time, and demonstrated that the uppermost ocean crust consisted of basaltic lavas. This achievement received a personal letter of congratulations from U.S. President Kennedy2. More than 50 years later, and since the launch of DSDP in 1968, oceanic basement has been drilled in a range of geodynamic settings. The first intentional drilling efforts in the oceanic crust were DSDP Leg 34 in 1973–1974, on the Nazca Plate in the Eastern Pacific Ocean (Yeats et al., 1976), and DSDP Leg 37 in 1974 on the western flank of the Mid-Atlantic Ridge, south of the Azores Plateau (Aumento et al., 1977). These Legs recovered for the first time substantial cores of basalt from the upper oceanic crust. In the Atlantic, cores from Site 334 also recovered small amounts of gabbros and serpentinized peridotites for the first time at shallow subseafloor depths, demonstrating, together with future seafloor geology, marine geophysics, and drilling along the Mid-Atlantic Ridge, that the layered “Penrose” model from the early 1970s (Anonymous, 1972) is not applicable worldwide. Since then, scientific ocean drilling has significantly contributed to our understanding of the variability of the architecture of oceanic lithosphere. The style of accretion critically depends on the balance between magma production, hydrothermal cooling, and tectonics, which is on the first order related to spreading rate. Seismic, bathymetric, and marine geological observations indicate that ocean crust formed at fast spreading rates (full rate >80 mm per year) is much less variable than crust formed at slow spreading rates (<40 mm per year) and is closer to the ideal “Penrose” layered model developed from ophiolites (Anonymous, 1972). Yet, numerous basic, direct observations regarding the architecture of in situ, present-day ocean crust, including rock types, deformation, or geochemistry, are yet to be made. A compilation of holes into the ocean crust cored by scientific ocean drilling since the beginning of DSDP is presented in Table 4.2.1.1 and Figures 4.2.1.1 and 4.2.1.2. Only 37 holes deeper than 100 m have been cored in oceanic crust since DSDP Leg 37 in 1974 (Figure 4.2.1.2). The recovered material represents <2% of the cores recovered to date by DSDP, ODP, and IODP. In spite of this cursory sampling, scientific drilling has contributed significantly to advancing knowledge of ocean crust architecture, mid-ocean-ridge accretion, and hydrothermal processes (e.g., Alt et al., 1996; Blackman et al., 2011; Dick et al., 2000; Dick, Natland, & Ildefonse, 2006; Ildefonse et al., 2007a,b; Ildefonse, Rona, & Blackman, 2007c; Teagle, Alt, & Halliday, 1998; Wilson et al., 2006). 2. http://nationalacademies.org/includes/KennedyTelegram.pdf.
TABLE 4.2.1.1 Drill Holes into Oceanic Basement in Intact Crust and Tectonically Exposed Lower Crust and Upper Mantle. This Compilation Does Not Include Other “Hard Rock” Drill Holes in Oceanic Plateaus, Arc Basement, Hydrothermal Mounds, or Passive Margins. Age (my)
Sediment Thickness (m)
Basement Penetration (m)
Recovery (%)
Spreading Rate
Lithology
Comments
11°09.21′S 70°31.56′E
Indian
2844.5
30
506
81
50
S/I
Basaltic lavas
Projection of Chagos-Laccadive Plateau
257
30°59.16′S 108°20.99′E
Indian
5278
120
262
65
50
S/I
Basalt and breccia
Wharton Basin off Perth, Australia
34
319A
13°01.04′S 101°31.46′W
Pacific
4296
16
98
59
25
F
Basaltic lavas
Bauer Deep, 13°S East Pacific Rise
37
332A
36°52.72′N 33°38.46′W
Atlantic
1851
3.5
104
331
10
S
Basalt, basalt breccia, and interlayered sediments
Mid-Atlantic Ridge 36°–37°N
37
332B
36°52.72′N 33°38.46′W
Atlantic
1983
3.5
149
583
21
S
Basalt and Mid-Atlantic basalt breccia Ridge 36°–37°N
37
333A
36°50.45′N 33°40.05′W
Atlantic
1665.8
3.5
218
311
8
S
Basalt and Mid-Atlantic basalt breccia Ridge 36°–37°N
37
335
37°17.74′N 35°11.92′W
Atlantic
3188
15
454
108
38
S
Basaltic lavas
Mid-Atlantic Ridge 36°–37°N
45
395A
22°45.35′N 46°04.90′W
Atlantic
4485
7.3
92
571
18
S
Basaltic lavas and breccia
Mid-Atlantic Ridge 23°N
Hole
Location
24
238
26
Formation and Evolution of Oceanic Lithosphere Chapter | 4.2.1 451
Ocean
Water Depth (m)
Leg/ Expedition
Continued
TABLE 4.2.1.1 Drill Holes into Oceanic Basement in Intact Crust and Tectonically Exposed Lower Crust and Upper Mantle. This Compilation Does Not Include Other “Hard Rock” Drill Holes in Oceanic Plateaus, Arc Basement, Hydrothermal Mounds, or Passive Margins.—cont’d
Ocean
Water Depth (m)
Age (my)
Sediment Thickness (m)
Basement Penetration (m)
Recovery (%)
Spreading Rate
Lithology
Comments
22°58.88′N 43°30.95′W
Atlantic
4450
9
126
96
33
S
Basaltic lavas
Mid-Atlantic Ridge 23°N
396B
22°59.14′N 43°30.90′W
Atlantic
4459
13
151
255
23
S
Basalt and breccia
Mid-Atlantic Ridge 23°N
49
410A
45°30.53′N 29°28.56′W
Atlantic
2987
9
331
49
38
S
Basaltic lavas
Mid-Atlantic Ridge 45°N
49
412A
36°33.74′N 33°09.96′W
Atlantic
2626
1.6
163
131
18
S
Basalt flows and intercalating limestone
Mid-Atlantic Ridge 33°N
51–53
417A
25°06.63′N 68°02.48′W
Atlantic
5478.2
110
208
209
61
S
Basaltic lavas
Western Atlantic
51–53
417D
25°06.69′N 68°02.81′W
Atlantic
5489
110
343
366
70
S
Basaltic lavas
Western Atlantic
51–53
418A
25°02.10′N 68°03.44′W
Atlantic
5519
110
324
544
72
S
Basaltic lavas
Western Atlantic
54
428A
09°02.77′N 105°26.14′W
Pacific
3358.5
2.3
63
53
39
F
Basaltic lavas
9°N East Pacific Rise
63
469
32°37.00′N 120°32.90′W
Pacific
3802.5
17
391
63
34
I
Basaltic lavas
Off California Coast
Leg/ Expedition
Hole
Location
45
396
46
63
470A
28°54.46′N 117°31.11′W
Pacific
3554.5
15
167
49
33
I
Basaltic lavas
Off California Coast
65
482B
22°47.38′N 107°59.60′W
Pacific
3015
0.5
137
93
54
I
Massive basalt and interlayered sediment
Off Gulf of California
65
482D
22°47.31′N 107°59.51′W
Pacific
3015
0.5
138
50
50
I
Massive basalt and interlayered sediment
Off Gulf of California
65
483
22°53.00′N 108°44.90′W
Pacific
3084
2
110
95
40
I
Massive basalt and pillow basalt with interlayered sediments
Off Gulf of California
65
483B
22°52.99′N 108°44.84′W
Pacific
3084
2
110
157
47
I
Massive basalt and pillow basalt with interlayered sediments
Off Gulf of California
65
485A
22°44.92′N 107°54.23′W
Pacific
2996.5
1.2
153.5
178
51
I
Massive basalt and interlayered sediments
Off Gulf of California
Continued
TABLE 4.2.1.1 Drill Holes into Oceanic Basement in Intact Crust and Tectonically Exposed Lower Crust and Upper Mantle. This Compilation Does Not Include Other “Hard Rock” Drill Holes in Oceanic Plateaus, Arc Basement, Hydrothermal Mounds, or Passive Margins.—cont’d
Ocean
Water Depth (m)
Age (my)
Sediment Thickness (m)
Basement Penetration (m)
Recovery (%)
Spreading Rate
Lithology
Comments
1°13.63′N 83°44.06′W
Pacific
3466.9
6.6
264
73
60
I
Basaltic lavas
South Flank of Costa Rica Rift
504B
1°13.611′N 83°43.818′W
Pacific
3474
6.6
270
1841
20
I
Basalt, stockwork, and diabase
South Flank of Costa Rica Rift
82
559
35°07.45′N 40°55.00′W
Atlantic
3754
35
238
63
37
S
Basaltic lavas
West Flank of MidAtlantic Ridge 35°N
82
562
33°08.49′N 41°40.76′W
Atlantic
3172
12
240
90
45
S
Pillow basalt and massive basalt
West Flank of MidAtlantic Ridge 33°N
82
564
33°44.36′N 43°46.03′W
Atlantic
3820
35
284
81
43
S
Pillow basalt and minor massive basalt
West Flank of MidAtlantic Ridge 34°N
91
595B
23°49.34′S 165°31.61′W
Pacific
5615
80
70
54
28
F
Vesicular Central South Pacific aphyric basalt
92
597C
18°48.43′S 129°46.22′W
Pacific
4164
30
53
91
53
F
Massive basalt flows
Leg/ Expedition
Hole
Location
68
501
69/70/83/111/ 137/140/ 148
West Flank South East Pacific Rise 18°S
106/109
648B
22°55.32′N 44°56.825′W
Atlantic
3326
0
0
50
12
S
Pillow basalt
Mid-Atlantic Ridge 23°N
109
670A
23°9.996′N 45°1.932′W
Atlantic
3625
0
0
77
6
S
Serpentinized peridotite
Mid-Atlantic Ridge, MARK, 23°N
118/176
735B
32°43.395′S 57°15.959′E
Indian
720
11.8
0
1508
86
S
Gabbro
Atlantis Bank, Southwest Indian Ridge
123
765D
15°58.56′S 117°34.51′E
Indian
5713.8
140
928
267
31
F
North–east mid-oceanridge basaltic lavas
Argo Abyssal Plain
129/185
801C
18°38.538′N 156°21.59′E
Pacific
5674
170
462
414
47
F
Pillow basalt, basalt flows, and breccias
Western North Pacific
129
802A
12°5.778′N 153°12.63′E
Pacific
5980
120
509
51
33
F
Basaltic lavas
Western North Pacific
136
843B
19°20.54′N 159°5.68′W
Pacific
4418
95
243
71
37
F
Basaltic lavas
West of Hawaii
147
894E
2°18.059′N 101°31.524′W
Pacific
3014
1
0
29
11
F
Gabbro
Hess Deep
147
894F
2°17.976′N 101°31.554′W
Pacific
3025
1
0
26
7
F
Gabbro
Hess Deep
147
894G
2°17.976′N 101°31.554′W
Pacific
3023
1
0
127.5
35
F
Gabbro
Hess Deep
Continued
TABLE 4.2.1.1 Drill Holes into Oceanic Basement in Intact Crust and Tectonically Exposed Lower Crust and Upper Mantle. This Compilation Does Not Include Other “Hard Rock” Drill Holes in Oceanic Plateaus, Arc Basement, Hydrothermal Mounds, or Passive Margins.—cont’d
Ocean
Water Depth (m)
Age (my)
Sediment Thickness (m)
Basement Penetration (m)
Recovery (%)
Spreading Rate
2°16.638′N 101°26.766′W
Pacific
3821
1
0
17
14
895B
2°16.638′N 101°26.760′W
Pacific
3821
1
0
10
147
895C
2°16.632′N 101°26.772′W
Pacific
3820
1
0
147
895D
2°16.638′N 101°26.778′W
Pacific
3821
1
147
895E
2°16.788′N 101°26.790′W
Pacific
3753
147
895F
2°16.902′N 101°26.790′W
Pacific
148
896A
1°13.006′N 83°43.392′W
153
920B
153 153
Leg/ Expedition
Hole
Location
147
895A
147
Lithology
Comments
F
Serpentinized peridotite
Hess Deep
10
F
Serpentinized peridotite
Hess Deep
38
15
F
Serpentinized peridotite
Hess Deep
0
94
20
F
Serpentinized peridotite
Hess Deep
1
0
88
37
F
Serpentinized peridotite
Hess Deep
3693
1
0
26
8
F
Serpentinized peridotite
Hess Deep
Pacific
3459
6.6
179
290
27
I
Basaltic lavas
South Flank of Costa Rica Rift
23°20.310′N 45°1.038′W
Atlantic
3339
<1
0
126.4
35.3
S
Serpentinized peridotite
Mid-Atlantic Ridge, MARK, 23°N
920D
23°20.322′N 45°1.044′W
Atlantic
3338
<1
0
200.8
47.3
S
Serpentinized peridotite
Mid-Atlantic Ridge, MARK, 23°N
921A
23°32.460′N 45°1.866′W
Atlantic
2488
<1
0
17.1
18.1
S
Gabbro
Mid-Atlantic Ridge, MARK, 23°N
153
921B
23°32.478′N 45°1.842′W
Atlantic
2490
<1
0
44.1
19.4
S
Gabbro
Mid-Atlantic Ridge, MARK, 23°N
153
921C
23°32.472′N 45°1.830′W
Atlantic
2495
<1
0
53.4
11.4
S
Gabbro
Mid-Atlantic Ridge, MARK, 23°N
153
921D
23°32.442′N 45°1.830′W
Atlantic
2514
<1
0
48.6
12.7
S
Gabbro
Mid-Atlantic Ridge, MARK, 23°N
153
921E
23°32.328′N 45°1.878′W
Atlantic
2456
<1
0
82.6
21.4
S
Gabbro
Mid-Atlantic Ridge, MARK, 23°N
153
922A
23°33.162′N 45°1.926′W
Atlantic
2612
<1
0
14.6
63.2
S
Gabbro
Mid-Atlantic Ridge, MARK, 23°N
153
922B
23°31.368′N 45°1.926′W
Atlantic
2612
<1
0
37.4
25.6
S
Gabbro
Mid-Atlantic Ridge, MARK, 23°N
153
923A
23°32.556′N 45°1.896′W
Atlantic
2440
<1
0
70
57.2
S
Gabbro
Mid-Atlantic Ridge, MARK, 23°N
153
924B
23°32.460′N 45°0.858′W
Atlantic
3170
<1
0
30.8
8.7
S
Gabbro
Mid-Atlantic Ridge, MARK, 23°N
153
924C
23°32.496′N 45°0.864′W
Atlantic
3177
<1
0
48.5
23.1
S
Gabbro
Mid-Atlantic Ridge, MARK, 23°N
168
1025C
47°53.250′N 128°38.880′W
Pacific
2602
1.237
106
41
37
I
Basalt
Juan de Fuca Flank
168
1026B
47°45.759′N 127°45.552′W
Pacific
2658
3.511
256
39
5
I
Basalt
Juan de Fuca Flank
168
1026C
47°46.261′N 127°45.186′W
Pacific
2669
3.516
229
19
3.5
I
Basalt
Juan de Fuca Flank
Continued
TABLE 4.2.1.1 Drill Holes into Oceanic Basement in Intact Crust and Tectonically Exposed Lower Crust and Upper Mantle. This Compilation Does Not Include Other “Hard Rock” Drill Holes in Oceanic Plateaus, Arc Basement, Hydrothermal Mounds, or Passive Margins.—cont’d
Ocean
Water Depth (m)
Age (my)
Sediment Thickness (m)
Basement Penetration (m)
Recovery (%)
Spreading Rate
Lithology
Comments
47°46.776′N 128°7.320′W
Pacific
2645
2.621
290
48
6.5
I
Basalt
Juan de Fuca Flank
1105A
32°43.135′S 57°16.652′E
Indian
714
11.8
0
158
75
S
Gabbro
Atlantis Bank, Southwest Indian Ridge
185
1149D
31°18.79′N 143°24.03′E
Pacific
5818
133
307
133
17
F
Pillow basalt, basalt flows, and breccias
Western North Pacific
187
1162B
44°37.9′S 129°11.3′′E
Indian
5464
18
333
59
17
I
Basaltic lavas and breccia
Australian-Antarctic Discordance
187
1163A
44°25.5′S 126°54.5′E
Indian
4354
17
161
47
33
I
Basaltic lavas
Australian-Antarctic Discordance
187
1164B
43°45.0′S 127°44.8′E
Indian
4798
18.5
150
66
16
I
Basaltic lavas
Australian-Antarctic Discordance
191
1179D
41°04.8′N 159°57.8′E
Pacific
5563.9
129
377
98
44
F
Basaltic lavas
Western North Pacific
200
1224F
27°53.36′N 141°58.77′W
Pacific
4967.1
46
28
147
26
F
Basaltic lavas
Central Pacific
203
1243B
5°18.07′N 110°04.58′W
Pacific
3868
11
110
87
25
F
Basaltic lavas
Western Flank East Pacific Rise 5°N
Leg/ Expedition
Hole
Location
168
1032A
179
206
1256C
6°44.18′N 91°56.06′W
Pacific
3634.7
15
251
89
61
F
Basaltic lavas
Cocos plate eastern Flank East Pacific Rise
206/ 309/312
1256D
6°44.16′N 91°56.06′W
Pacific
3634.7
15
250
1257.1
37.1
F
Basaltic Cocos plate Eastern lavas, sheeted Flank East Pacific dike, and Rise varitextured gabbro
209
1268A
14°50.755′N 45°4.641′W
Atlantic
3007
0
147.6
53.3
S
Serpentinized peridotite
Mid-Atlantic Ridge 15°20′N
209
1270A
14°43.342′N 44°53.321′W
Atlantic
1951
0
26.9
12.2
S
Gabbro
Mid-Atlantic Ridge 15°20′N
209
1270B
14°43.265′N 44°53.225′W
Atlantic
1909
0
45.9
37.4
S
Gabbro
Mid-Atlantic Ridge 15°20′N
209
1270C
14°43.284′N 44°53.091′W
Atlantic
1822
0
18.6
10.6
S
Gabbro
Mid-Atlantic Ridge 15°20′N
209
1270D
14°43.270′N 44°53.084′W
Atlantic
1817
0
57.3
13.4
S
Gabbro
Mid-Atlantic Ridge 15°20′N
209
1271A
15°2.222′N 44°56.887′W
Atlantic
3612
0
44.8
12.9
S
Serpentinized peridotite
Mid-Atlantic Ridge 15°20′N
209
1271B
15°2.189′N 44°56.912′W
Atlantic
3585
0
103.8
15.3
S
Serpentinized peridotite
Mid-Atlantic Ridge 15°20′N
209
1272A
15°5.666′N 44°58.300′W
Atlantic
2560
0
131
28.6
S
Serpentinized peridotite
Mid-Atlantic Ridge 15°20′N Continued
TABLE 4.2.1.1 Drill Holes into Oceanic Basement in Intact Crust and Tectonically Exposed Lower Crust and Upper Mantle. This Compilation Does Not Include Other “Hard Rock” Drill Holes in Oceanic Plateaus, Arc Basement, Hydrothermal Mounds, or Passive Margins.—cont’d Leg/ Expedition
Hole
Location
Ocean
Water Depth (m)
Sediment Thickness (m)
Basement Penetration (m)
Recovery (%)
Spreading Rate
209
1274A
15°38.867′N 46°40.582′W
Atlantic
3940
0
155.8
22.2
209
1275B
15°44.486′N 46°54.208′W
Atlantic
1562
0
108.7
209
1275D
15°44.440′N 46°54.217′W
Atlantic
1554
0
301
U1301A 47°45.209′N 127°45.833′W
Pacific
2656
3.5
301
U1301B 47°45.229′N 127°45.826′W
Pacific
2655
304
U1309B
30°10.108′N 42°7.110′W
Atlantic
304/ 305
U1309D 30°10.120′N 42°7.113′W
Atlantic
Age (my)
Lithology
Comments
S
Serpentinized peridotite
Mid-Atlantic Ridge 15°20′N
43.1
S
Gabbro
Mid-Atlantic Ridge 15°20′N
209
50
S
Gabbro
Mid-Atlantic Ridge 15°20′N
262
108
0
I
Basaltic lavas
Juan de Fuca Ridge Flank; no coring (CORK)
3.5
265
318
12.9
I
Basaltic lavas
Juan de Fuca Ridge Flank; recovery is only for the 232 m of cored basement
1642
2
2
99.8
45.9
S
Gabbro
Mid-Atlantic Ridge 30°N
1645
2
2
1413.3
74.8
S
Gabbro
Mid-Atlantic Ridge 30°N
327
U1362A 47°45.663′N 127°45.672′W
Pacific
2661
3.5
236
292
29.6
I
Basaltic lavas
Juan de Fuca Ridge Flank; recovery is only for the 150 m of cored basement
336
U1383C 22°48.124′N 46°03.166′W
Atlantic
4414
8
38.3
293.2
26.5
S
Basaltic lavas
North Pond. Corked Hole. Upper 31.2 m of basement not cored
345
U14515I 2°15.155′N 101°32.662′W
Pacific
4841.8
1
2
35.2
20
F
Gabbro
Hess Deep lower crust
345
U1415J
2°15.160′N Pacific 101° 32.662′W
4838.8
1
2
111.8
15.6
F
Gabbro
Hess Deep lower crust. Upper 15 m of basement not cored
345
U1415P 2°15.149′N 101°32.610′W
4852.7
1
2
107.9
32
F
Gabbro
Hess Deep lower crust. Upper 12.5 m of basement not cored
Pacific
S = Slow (<40 mm per year); I = Intermediate; F = Fast (>80 mm per year); MARK = Mid-Atlantic Ridge Kane Fracture Zone; CORK = Circulation Obviation Retrofit Kit.
462 Earth and Life Processes Discovered from Subseafloor Environments +ROHV'ULOOHGLQ2FHDQ&UXVW '6'3 2'3 ,2'3
;
; 8
; 8
$JHRI2FHDQLF&UXVW 0LOOLRQ
FIGURE 4.2.1.1 DSDP, ODP, and IODP holes drilled in ocean crust >100 mbsf from 1974 to 2013. Sites mentioned in the text are labeled. Also indicated are the locations of the four IODP expeditions reviewed herein. The seafloor age is based on age grid by Müller et al. (2008, revised version 3; www.earthbyte.org/). This map does not include “hard rock” drill holes in oceanic plateaus, arc basement, hydrothermal mounds, or passive margins.
Hole 504B, located on 6.9 my crust formed at an intermediate rate at the Costa Rica Rift (Figure 4.2.1.1), remains the second deepest hole (2111 meters below seafloor (mbsf)) in all of scientific ocean drilling and the deepest into oceanic basement (1836 m subbasement; Alt et al., 1996). This site was host to drilling and other experiments over eight DSDP and ODP cruises, and was the first hole to penetrate completely through the volcanic lava sequences and ∼1 km into sheeted dikes in intact, layered oceanic crust. It remains a reference hole for hydrothermal alteration of the upper ocean crust (e.g., Alt, Honnorez, Laverne, & Emmermann, 1986; Alt, Muehlenbachs, & Honnorez, 1986), and for the geological significance of seismic Layers 2A and 2B (e.g., Carlson, 2011), thought to represent a porous lava flow layer and the underlying sheeted dike complex, respectively (e.g., Herron, 1982). In fast-spread crust, the Layer 2/3 boundary is classically assumed to be the boundary between the upper crust (sheeted dikes and lavas) and the underlying plutonic gabbroic crust. This transition between typical layer 2 and layer 3 seismic velocity gradients with depth has been sampled in situ Hole 504B, where it lies in the sheeted dike complex, and appears related to changes in alteration and porosity structure (Carlson, 2010; Detrick, Collins, Stephen, & Swift, 1994).
Formation and Evolution of Oceanic Lithosphere Chapter | 4.2.1 463
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FIGURE 4.2.1.2 Compilation chart showing holes drilled >100 m in intact crust and tectonically exposed lower crust and upper mantle from 1974 to 2013 (drill hole locations in Figure 4.2.1.1). For each hole are indicated the hole number and the recovery (in percent) for each lithology. This compilation does not include “hard rock” drill holes in oceanic plateaus, arc basement, hydrothermal mounds, or passive margins.
A complete, continuous section of intact, layered fast-spread crust down to the cumulate gabbro layers has yet to be drilled and remains a first-order scientific target of ocean drilling for the ocean crust research community (e.g., Chikyu+10 Steering Committee, 2013; Dick & Mével, 1996; Ildefonse et al., 2010a; I ldefonse et al., 2010b; Ildefonse, Christie, & the Mission Moho Workshop S teering Committee, 2007b; IODP S cience Plan 2013–2023, 2011; Ravelo et al., 2010; Teagle, Wilson, & Acton, 2004; Teagle et al., 2009). Over the last 10 years, Hole 1256D (ODP Leg 206, Integrated Ocean Drilling Program (IODP) Expeditions 309, 312, and 335; Figure 4.2.1.1) penetrated to the base of the sheeted dike complex and the uppermost gabbro in superfast-spread crust (i.e., >130 mm per year; Lonsdale, 1977); this was the first sampling of the transition to plutonic rocks in intact ocean crust (Teagle et al., 2006; Teagle, Ildefonse, Blum, & the Expedition 335 Scientists, 2012; Wilson et al., 2003; Wilson et al., 2006). An alternative to drilling deep through intact, fast-spread crust is to access the deepest part of the crust in tectonic windows. This strategy was recently used at Hess Deep, in the eastern equatorial Pacific just off the East Pacific Rise (EPR), during IODP Expedition 345 to access lower crustal gabbroic rocks (Gillis, Snow, Klaus, & the Expedition 345 Scientists, 2014; Gillis et al., 2014a; Figure 4.2.1.1).
464 Earth and Life Processes Discovered from Subseafloor Environments
In 1986, serpentinized mantle peridotites were intentionally drilled at a midocean ridge for the first time during ODP Leg 109 (Site 670; Detrick et al., 1988) on the west wall of the Mid-Atlantic Ridge median valley near 23°10′N. In the same area, south of the Kane Fracture Zone, 95 m of serpentinized peridotites were recovered from a 200-m-deep hole during ODP Leg 153 (Site 920; Cannat et al., 1995). Ten years later, ODP Leg 209 (Sites 1268–1275) returned to drill in the peridotite-rich area around the 15°20′N fracture zone on the MidAtlantic Ridge (Kelemen, Kikawa, Miller, & Shipboard Scientific Party, 2007). Thirteen holes at six sites along the spreading axis penetrated mantle peridotite and gabbroic rocks in proportions that were roughly 70/30, similar to what was previously sampled by dredging and submersible on the seafloor in the same area (compiled data and references in Kelemen et al., 2004). These findings, together with geophysics and seafloor geology, contributed to the understanding that slowspreading oceanic crust can have a very variable architecture, locally consisting of relatively small gabbro bodies intruded into serpentinized mantle peridotites, and that spreading in nonlayered, magma-poor crust is largely accommodated by detachment faulting (e.g., Cannat, 1993; Dick, 1989; Escartín et al., 2008; Escartín & Canales, 2011; Kelemen et al., 2007; Lagabrielle, Bideau, Cannat, Karson, & Mével, 1998). In this type of crust, the spatially and temporally variable interplay of magmatic and tectonic processes is locally expressed in the development of Oceanic Core Complexes (OCCs). OCCs are domal features that result from the exhumation of lower crustal and upper mantle rocks through long-lived (typically 1–3 my) detachment faulting (e.g., Cann et al., 1997; Dick, Tivey, & Tucholke, 2008; Escartín & Canales, 2011; Escartín, Mével, MacLeod, & McCaig, 2003; MacLeod et al., 2009; Smith, Escartín, Schouten, & Cann, 2008; Tucholke & Lin, 1994; Tucholke, Lin, & Kleinrock, 1998). The first deep hole drilled in an OCC was the 1508-m-deep ODP Hole 735B in gabbroic rocks of the Atlantis Bank on the Southwest Indian Ridge (Dick et al., 2000). More recently, IODP successfully sampled another OCC, with Hole U1309D penetrating 1415 m of gabbroic rocks at the Atlantis Massif on the Mid-Atlantic Ridge during Expeditions 304 and 305 (Blackman et al., 2011; Ildefonse et al., 2007a). This chapter draws from results of 10 years of IODP ocean drilling (2003– 2013) in igneous oceanic crust, focusing on the two deep holes in slow-spread (Hole U1309D) and fast-spread (Hole 1256D) crust, and briefly summarizing the outcome of the recent IODP Expedition 345 at Hess Deep. These drilling expeditions have improved our understanding of crustal architecture and accretion processes at mid-ocean ridges.
4.2.1.2 DEEP DRILLING IN SLOW-SPREAD CRUST: THE ATLANTIS MASSIF 4.2.1.2.1 Background IODP Expeditions 304 and 305, from November 2004 to March 2005, targeted one of the emblematic OCCs, the Atlantis Massif, at 30°N on the Mid-Atlantic
Formation and Evolution of Oceanic Lithosphere Chapter | 4.2.1 465
Ridge (Blackman, Cann, Janssen, & Smith, 1998; Blackman et al., 2002; Cann et al., 1997; Schroeder & John, 2004). The Atlantis Massif is a young (0.5–2 my) OCC, situated at the inside corner of the intersection between the Mid-Atlantic Ridge and the Atlantis Fracture zone at 30°N. Seafloor lithologies observed by dredging and submersible sampling include serpentinites, gabbros, and basaltic rocks (Figure 4.2.1.3; Cann et al., 1997; Blackman et al., 2002). The active, serpentinite-hosted Lost City hydrothermal vent field (Früh-Green et al., 2003; Kelley et al., 2001) is located immediately south of the summit of the massif. Localized deformation and fluid flow associated with detachment faulting are documented from seafloor sampling on the southern wall of the massif (Boschi, Früh-Green, Delacour, Karson, & Kelley, 2006; Karson et al., 2006; Schroeder & John, 2004). Two deep holes were drilled at Site U1309 during IODP Expeditions 304 and 305, in the footwall of the detachment fault (Figure 4.2.1.3; Blackman et al., 2006, 2011). In Û :
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ODQ W
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IODP Hole U1309D Seaflor samples : basalt gabbro serpentinized peridotite Lost City hydrothermal field
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'HSWK P FIGURE 4.2.1.3 Bathymetric map of Atlantis Massif showing geological and some seismic (reflection and refraction) data coverage prior to drilling, and the location of Site U1309 (yellow symbol (pale gray in print versions)), modified from Blackman et al. (2006). The WNW-ESE trending corrugations on the domal core mark the exposed detachment fault; hanging wall block(s) are volcanic.
466 Earth and Life Processes Discovered from Subseafloor Environments
addition, IODP Expedition 340T revisited this site in 2012 to conduct 5 days of borehole logging in Hole U1309D, to complete the logging program that was hampered by bad weather conditions at the end of IODP expedition 305 (Blackman et al., 2006). The program included completing sonic logging to the bottom of the hole, and acquiring new seismic imaging data from a Vertical Seismic Profile (VSP) experiment (Expedition 340T Scientists, 2012).
4.2.1.2.2 Revisiting the Geophysical Signature of the Atlantis Massif Based on the dominant occurrence of serpentinized mantle peridotites sampled on the south flank of the southern ridge (Figure 4.2.1.3; e.g., Blackman et al., 2002; Boschi et al., 2006), and on the predrilling interpretation of available seismic (Figure 4.2.1.3; Collins & Detrick, 1998; Collins, Tucholke, & Canales, 2002; Canales, Tucholke, & Collins, 2004) and gravimetric (Blackman, Cann, Janssen, & Smith, 1998; Nooner, Sasagawa, Blackman, & Zumberge, 2003) data, one of the goals of IODP Expeditions 304/305 was to drill and sample the core of an OCC that was inferred to be mostly made of mantle peridotites. It was hoped that fresh mantle peridotite was present at reasonably shallow depths (∼800 mbsf) accessible by conventional nonriser drilling and hence, could be sampled for the first time (see Blackman et al., 2011; for a summary of the predrilling rationale). The recovery of a 1415-m-long gabbroic section demonstrated that geological sampling by drilling is essential to better constrain the interpretation of geophysical data in ocean crust, particularly in slow-spread crust, which can be lithologically heterogeneous on small scales. Drill cores, together with wireline geophysical measurements provided critical observations for reexamining the precruise geophysical interpretations and for proposing a revised analysis of the lithological architecture of the Atlantis Massif and other OCCs (Blackman, Karner, & Searle, 2008; Blackman, Canales, & Harding, 2009; Blackman and Collins, 2010; Canales, Tucholke, Xu, Collins, & DuBois, 2008; Collins, Blackman, Harris, & Carlson, 2009; Henig, Blackman, Harding, Canales, & Kent, 2012). The positive 30–40 mGal residual gravity anomaly centered on the core of the OCC is now considered to reflect a gabbroic core with a density of 2900 kg/m3, the adjacent basaltic block is significantly fractured with an average density of 2600 kg/m3, and the portion of the lithosphere deeper than 1500 mbsf that has density lower than mantle rock (gabbroic, and/or significantly altered peridotite) is ∼3-km thick within the domal core (Blackman, Karner, & Searle, 2008). Postdrilling seafloor refraction modeling showed that shallow mantle velocities are not required (Figure 4.2.1.4; Collins et al., 2009) and velocities in the upper several hundred meters are typical of mafic rock (≤6.5 km/s). Tomographic inversion of refracted arrivals recorded by a 6-km multichannel seismic (MCS) streamer (Canales et al., 2008) produced a similar velocity structure to 1.0–1.5 km depth across the Central Dome (Figure 4.2.1.4(D)). Further detailed processing of the MCS data, continued downward using the “synthetic ocean bottom experiment” technique (Henig et al., 2012),
Formation and Evolution of Oceanic Lithosphere Chapter | 4.2.1 467
2
0.3
4
5
6
7
South
North
Southern Ridge
Central Dome
7
South
North
6 5 4
3
Hole U1309D
-15
(C) Depth (km)
3
0
-10
3
-5
0
5
2
-15
Hole U1309D
-10
2
8 7 6 5 4
1
3
10 km
-5
0
5 km
3 2
velocity (km/s)
Depth (km)
(B) 1
velocity (km/s) 2
velocity (km/s)
Depth (km)
(A)
Two-way travel time (s)
(D)
CMP FIGURE 4.2.1.4 Seismic imaging of the Atlantis Massif. (A) Preferred velocity model along the NOBEL refraction Line 9 on the central dome (location in Figure 4.2.1.3), determined from forward modeling by Collins et al. (2009). (B) Velocity model based on inversion of multichannel seismic (MCS) refractions for Line 4 (location in Figure 4.2.1.3) by Canales et al. (2008). (C) Velocity model based on inversion of MCS refractions for Line 4 by Henig et al. (2012). (D) Time-migrated section for MCS Line 4 (Canales et al., 2004); D: D-reflector; CMP: Common-Midpoint.
reveals significant spatial (both along-axis and along-flow line) variability of the P-wave velocity structure at the scale of the Atlantis Massif. This variability is generally consistent with lithologically heterogeneous crust, made of gabbroic plutons intruded into a dominantly peridotitic (serpentinized) crust. The western two-thirds of the southern ridge, to date unsampled by drilling, is dominated by
468 Earth and Life Processes Discovered from Subseafloor Environments
low-velocity material that is likely to be serpentinized peridotites (at least down to ∼1 km below seafloor), whereas the eastern part displays a higher velocity body that is interpreted as being a gabbro pluton. Analysis of a 40-km-long EW air gun refraction profile across the central dome provided constraints on velocities to depths as great as 7 km (Blackman and Collins, 2010). Traveltime tomography suggests that significant volumes of rock with mantle-like velocity (>7.5 km/s) only occur more than ∼5 km below seafloor within the dome, and only below 6 km in the axial valley. Initial results from IODP Expedition 340T seem to indicate that combined lithological changes and local increases in alteration could produce an impedance contrast strong enough to account for the reflectivity in the upper part of the section. In particular, the sonic log confirms that the borehole velocity of altered olivine-rich troctolite intervals in Hole U1309D may be sufficiently distinct from that of surrounding rocks (P-wave velocity (VP) ∼0.5 km/s slower) to produce an MCS reflection (Expedition 340T Scientists, 2012). However, the source(s) of the impedance contrast that gives rise to the strong reflection imaged throughout much of the dome (the “D-reflector”; Canales et al., 2004) is yet to be identified.
4.2.1.2.3 The Interplay between Magmatism and Tectonics Controls the Development of OCCs Hole U1309D sampled a 1415-m-long section of gabbroic rocks in the core of the Atlantis Massif OCC, with 75% recovery (Figure 4.2.1.5). This is the second longest core drilled in slow-spread crust, after ODP Hole 735B in the Southwest Indian Ocean (Dick et al., 2000). The borehole lithology at Site U1309D contrasts remarkably with the seafloor lithology of the Atlantis M assif. Although seafloor sampling dominantly recovered serpentinized peridotites on the south face of the Southern Ridge (approximately 70% of the recovered samples; Blackman et al., 2002; Karson et al., 2006; Schroeder & John, 2004), only three thin (<1 m) intervals of ultramafic rocks interpreted as residual mantle peridotites, intercalated within gabbros in the upper 225 m of the section (Tamura, Arai, Ishimaru, & Andal, 2008), were recovered in Holes U1309B and U1309D on the Central Dome, 5 km to the north. It is worth noting that including the completion of IODP Expeditions 304/305, ODP and IODP has drilled 16 holes in four different OCCs, and that all holes had recovered exclusively gabbroic sections (Ildefonse et al., 2007a). Hence OCCs are associated with significant magmatism. In contrast with early models (e.g., Escartin et al., 2003; Tucholke & Lin, 1994; Tucholke et al., 1998), this basic observation indicates that OCCs are not the magma-starved endmember of slow-spreading ridges. The revised model proposed that, in heterogeneous crust composed of gabbroic plutons intruded into variably serpentinized mantle peridotites (resulting from episodic magmatic activity), the development of an OCC is associated with, or results from a period of enhanced magmatism at depth. If this model is correct, the
(A)
(B)
0
0.5
1
(C)
Mg# (= Mg/[Mg + Fe]) (whole rock [molar])
20 m running average Fraction of cored interval
50 0
60
70
80
90
(D)
Plastic and magmatic foliation (intensity, 5 m ave.)
Cataclasis and veining (intensity, 5 m ave.)
0
0
1
2
2
4
(E)
Total alteration (%, 5 m intervals) 0
20
40
60
(F)
(G)
Bulk density (g/cc)
80 2.4 2.6 2.8
3
3.2
(H)
Vp (Km/s) 4
5
6
(I)
High-stability magnetization (inclination [°]) 7
-60
-30
0
Zircon Pb/U Age (my) 1.0
1.1
1.2
1.3
1.4
1.17 ± 0.02 my
400
800
my 1.24 ± 0.02 my
Depth (mbsf)
600
1000
1200
1400 Basalt/Diabase Oxide gabbro Gabbro, gabbronorite Troctolitic gabbro, olivine gabbro Troctolite Olivine-rich troctolite, dunite, harzburgite
Basalt/Diabase Oxide gabbro Gabbro, gabbronorite, olivine-bearing gabbro Olivine gabbro, troctolitic gabbro
Magmatic Plastic
Veining Cataclasis
Discrete samples Logging runs 1 & 3 Logging runs 1 & 3 5 m running average
Discrete samples Logging: 304 5m running 305 average 340T
Single measurement Multiple measurements Discrete samples
Felsic dike Oxide gabbro
Geocentric axial dipole
Troctolite Olivine-rich troctolite, dunite, harzburgite
my Fault zones inferred from cataclastic deformation of fault gouge in core
Leucocratic gabbro
FIGURE 4.2.1.5 Downhole plots of data for IODP Hole U1309D (modified from Ildefonse et al., 2006). (A) Lithology (20 m running average, white = no recovery); (B) Mg# (shipboard data); (C) plastic and magmatic foliation intensity (shipboard macroscopic description); (D) cataclasis and veining intensity (shipboard macroscopic description); (E) total alteration (shipboard macroscopic description); (F) bulk density (from borehole and discrete sample shipboard measurements); (G) compressional velocity (from borehole and discrete sample shipboard measurements); (H) inclination of stable remanent magnetization (shipboard measurements); (I) ages obtained for core samples from zircon (postcruise data, Grimes et al., 2008. The gray vertical bars are interpretative coherent age blocks, above and below 600 mbsf, with the average ages indicated aside).
Formation and Evolution of Oceanic Lithosphere Chapter | 4.2.1 469
200
470 Earth and Life Processes Discovered from Subseafloor Environments
domal morphology of OCCs would essentially be the consequence of having larger gabbroic plutons being exhumed and unroofed by the associated detachment fault (Blackman et al., 2011; Ildefonse et al., 2007a). The model implies that serpentinized fault zones envelop gabbro bodies, explaining why serpentinites can be the dominant lithology on the top or flanks of OCCs, while drilling to date suggests that their core is dominantly gabbroic. The inference that OCCs are associated with significant magma emplacement in the crust is consistent with a series of recent numerical models, which suggest that the optimal magmatic input for the development of long-living detachment faults and OCCs corresponds to about half of the total spreading (Behn & Ito, 2008; Buck, Lavier, & Poliakov, 2005; Olive, Behn, & Tucholke, 2010; Tucholke, Behn, Buck, & Lin, 2008). It is also consistent with observations from the southwest Indian Ocean (Cannat et al., 2006), where the true amagmatic stage of spreading at the Southwest Indian Ridge corresponds to the development of a smooth topography through conjugate, long-lived detachment faulting, marked by broad ridges and the absence of fault scarps and volcanics. OCCs there are identified through their corrugated topography, and are interpreted to represent the intermediate level of magmatic input between the fully magmatic volcanic seafloor and the amagmatic smooth seafloor.
4.2.1.2.4 Paleomagnetic Data and Borehole Imaging Constrain Tectonic Rotation in OCCs Competing models for the development of OCCs through long-lived detachment faulting involve either displacement on planar, low-angle normal faults (e.g., Wernicke, 1995) or progressive shallowing by rotation of initially steeply dipping (>45°) faults as a result of flexural unloading (i.e., the “rolling-hinge” model; e.g., Lavier, Buck, & Poliakov, 1999). Magmatic rocks acquire a magnetization in the direction of the ancient geomagnetic field at the time of their cooling and crystallization that may be used as a marker for subsequent tectonic rotation. The different models of OCC development can be tested by comparing the orientation of paleomagnetic remanence vectors in recovered magmatic rocks to the expected orientation of the Earth’s magnetic field at the sampling location. However, a fundamental difficulty inherent in scientific ocean drilling is that core samples lack azimuthal orientation in the geographical reference frame. Hence, documenting the amount and style of tectonic rotation of OCC footwall sections using paleomagnetic data becomes a problem with too many unknowns and multiple solutions. The way around this problem is to combine paleomagnetic data with structural observations in the core and borehole wireline imaging (e.g., Haggas, Brewer, Harvey, & MacLeod, 2005; MacLeod, Parson, Sager, & the ODP Leg 135 Scientific Party, 1992). In the absence of borehole imaging, Garcés and Gee (2007) had to make an assumption about the orientation of the rotation axis for the proposed tectonic rotations inferred from paleomagnetic measurements in serpentinite and gabbro sections exhumed by
Formation and Evolution of Oceanic Lithosphere Chapter | 4.2.1 471
detachment faulting in the 15°20′N fracture zone region of the Mid-Atlantic Ridge, sampled during ODP Leg 209 (Kelemen et al., 2007). For the first time, Hole U1309D offered complete, high-quality Formation MicroScanner (FMS) imaging of the borehole coupled with high core recovery, enabling a comprehensive core-log integration study (Figure 4.2.1.6). Using paleomagnetic remanence data and FMS images to reconstruct the true orientation of core pieces, Morris et al. (2009) demonstrated that the upper 400 m of the footwall of the detachment fault in the Atlantis Massif experienced an ∼45° anticlockwise rotation around a ridge-parallel horizontal axis trending NNE (010°), providing unequivocal validation of the rolling-hinge model (e.g., Lavier et al., 1999) for the development of OCCs. Analysis of the high-quality FMS imagery from Hole U1309D also allowed the distribution of structures within the Atlantis Massif OCC footwall to be comprehensively documented (Pressling, Morris, John, & MacLeod, 2012), revealing details of the internal structure of an OCC
IODP core “X”
(B)
True N
Discrete samples
(A)
(C)
011° ± 6°
N
46° ± 6° CCW
Archive half core
Mean remanence (discrete samples) Reference vector
98 - 367 mbsf IODP core coordinates
True geographic coordinates
FIGURE 4.2.1.6 Effect of FMS-based core reorientation on paleomagnetic data from Hole U1309D, and resulting constraints on footwall rotation in the Atlantis Massif Oceanic Core Complex (OCC) (modified after Morris et al., 2009). (A) Paleomagnetic remanence directions from discrete samples (red symbols (dark gray in print versions)) and archive half core sections (green symbols (light gray in print versions)) in the unoriented IODP core reference frame have a scattered distribution as a result of the drilling process. (B) The same data after individual core pieces have been reoriented to a true geographic reference frame by matching structures observed in the core with the corresponding structures in the borehole wall (the orientation of which may be determined from the FMS imagery). The independently reoriented paleomagnetic data now display a clustered distribution from which a reliable mean remanence vector may be calculated. (C) Comparison of the full mean remanence vector from discrete samples with the reference geocentric axial dipole vector for Atlantis Massif allows the orientation of the axis of footwall rotation (tilting) and angle of rotation to be determined, indicating substantial rotation around a Mid-Atlantic Ridge (MAR)-parallel axis consistent with rolling-hinge models for OCC development.
472 Earth and Life Processes Discovered from Subseafloor Environments
for the first time. A zone of brittle fracturing dominates the upper 385 m of the hole, and structures indicating detachment-related damage zones extend down to 750 mbsf into the OCC footwall (See also Michibayashi et al., 2008).
4.2.1.2.5 Protracted Construction of the Lower Crust Cores recovered in the upper 120 m at Site U1309 contain 35–45% of diabase/basalt in the form of subhorizontal, 5- to 10-m-thick intrusive bodies (Blackman et al., 2011) with chilled margins against cataclastic gabbros, amphibole-rich breccias, and undeformed diabase. They are interpreted as dikes intruded into the detachment fault prior to tectonic rotation (McCaig & Harris, 2012). The composition of all diabase recovered at Site U1309D falls within the range of basalt compositions for the Mid-Atlantic Ridge in the 30°N region (Godard et al., 2009). The dominantly gabbroic sequence in Hole U1309D comprises hundreds of individual lithologic units ranging in composition from oxide-gabbro to olivine-rich troctolite (Figures 4.2.1.5 and 4.2.1.7), identified on the basis of modal composition, grain-size changes, and/or textural changes. Units vary in thickness, from centimeters to tens of meters (John et al., 2009). Many intrusive contacts were recovered and indicate a general relationship of more evolved rock types intruded into more primitive rock types, although the converse sense of intrusion is also observed. Leucocratic intrusions (centimeter-scale thickness) cut the sequence and, together with oxide-gabbro intervals, host zircon grains that have been used to obtain crystallization ages of the recovered crustal section. Ages range from 1.08 ± 0.07 my to 1.28 ± 0.05 my, which indicates that the magmatic accretion of the multiple units described above has occurred over a minimum of approximately 100–200 ky (Figure 4.2.1.5; Grimes, John, Cheadle, & Wooden, 2008). The distribution of ages is inconsistent with models requiring constant depth melt intrusion beneath the detachment fault. Instead, data suggest a protracted construction model (Grimes et al., 2008), in which gabbroic rocks intrude at various depths below the ridge axis over a depth range of at least 1.4 km, before being exhumed by the detachment fault. Modal mineralogy and bulk rock geochemistry (e.g., Figure 4.2.1.7(A) and (B)) of the gabbroic sequence are typical of a cumulate series crystallizing from a mid-ocean ridge basalt (MORB) source. However, the bulk composition of Hole U1309D does not represent the complement to basalts recently erupted at the nearby spreading axis (Godard et al., 2009). The composition of the recovered rocks spans a large spectrum (Figure 4.2.1.7), with some troctolites and gabbros, respectively having some of the highest and lowest MgO contents of oceanic crustal rocks sampled to date. The bulk trace element and major element chemistry are consistent with most or all of the section being crystallized from a MORB melt at depth. When all lithologies, including leucocratic intervals, are accounted for, an extra basaltic counterpart to this section is not required (Figure 4.2.1.8; see detailed discussion in Godard et al., 2009).
Formation and Evolution of Oceanic Lithosphere Chapter | 4.2.1 473
(C)
(A) Ol-rich troctolite Troctolite
Gabbro
Gabbro and Ol-gabbro Other Exp 304-305 samples
Intrusive contact Troctolite
(B)
Obscured contact
Reaction zone
Oxide gabbro
(D) Ol-rich troctolite
Atlantis massif (Southern wall)
Troctolite
Gabbro (SWIR)
Olivine gabbro
Gabbro (MAR)
Gabbro
Gabbro & troctolite (MAR-ODP Site 1275)
Gabbronorite Oxide gabbro
Basalt (MAR)
Leucocratic dike
Peridotite (MAR-15°20’N)
Basalt Diabase Peridotite
Peridotite (MARK) Impregnated Peridotite (MAR-15°20’N)
Oxide gabbro
Gabbro 2 cm
FIGURE 4.2.1.7 Petrology, geochemistry, and photographs of Site U1309 core samples (From Blackman et al., 2011). (A) Compositions of representative suite of samples from site U1309 are shown by small gray symbols. Other symbols indicate samples from within the intervals that are dominantly olivine-rich troctolite that were studied in detail by Drouin et al. (2009). (B) Bulk rock chemistry for Site U1309 and comparison with other areas (Godard et al., 2009). Note how troctolites and olivine-rich troctolites fill in previously sparsely populated Ni-Mg# space. (C) Photo of Core 304-U1309D-69R-1 shows intrusive contact between gabbro and troctolite and lower contact that is a reaction zone between the troctolite and a later oxide gabbro injection. (D) Oxide gabbro dike intrudes gabbro (Hole U1309D, 386 mbsf) with a sharp lower contact with gabbro (Grimes et al., 2008).
474 Earth and Life Processes Discovered from Subseafloor Environments
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FIGURE 4.2.1.8 Bulk composition of Hole U1309D compared to the compositions of bulk gabbroic crust at Hole 735B and of primary mid-ocean ridge basalt (MORB) magma illustrated on (A) Ca# versus Mg# and (B) extended trace element diagrams (modified from Godard et al., 2009). The bulk composition of Hole U1309D is calculated taking into account all gabbroic rocks (BH) and without the olivine-rich troctolites (BG). The estimate of the bulk composition of Hole U1309 (BH) is characterized by high Mg# compared to primary magma compositions. These high Mg# may indicate a cumulate composition complementary to that of MORB, but also an overestimation of the mass of the most primitive endmember of the gabbroic suite, or an underestimation of the mass of its most evolved endmembers in our calculations. These hypotheses are tested using two models: (1) BH + LD: a 95/5% mix of BH and leucocratic dyke (LD); (2) BG + LD: a 95/5% mix of BG and LD. Estimates of the bulk composition of Hole 735B (Hart et al., 1999; Dick et al., 2000) as well as the mean compositions of basalts at Site U1309 and Hole 735B are shown for comparison. Hole 735B “protolith estimate” was determined on the basis of analyses conducted on a small suite of fresh and unaltered samples representative of the “characteristic lithologies”
Formation and Evolution of Oceanic Lithosphere Chapter | 4.2.1 475
Relatively low-pressure crystallization depths (≤200 MPa, ∼6 km) are inferred based on the modal relationships and chemistry of cumulus and intercumulus phases (Godard et al., 2009; Suhr, Hellebrand, Johnson, & Brunelli, 2008). This conclusion is supported by a combination of rock ages, assumptions about fault geometry, and isotherm depth that constrain crystallization depths at 5–7 km (Grimes, Cheadle, John, Reiners, & Wooden, 2011).
4.2.1.2.6 Olivine-Rich Troctolites: Mantle Contribution to the Igneous Lower Crust Olivine-rich troctolites (>70% olivine) are present as relatively thin (approximately 1–12 m) units within two main intervals in Hole U1309D (approximately 310–350 mbsf and 1090–1235 mbsf; Figure 4.2.1.5). Bulk rock geochemical signatures of these troctolites are more primitive (high MgO) than other mafic samples from the ocean crust (Figure 4.2.1.7; Drouin, Godard, Ildefonse, Bruguier, & Garrido, 2009; Godard et al., 2009). They display classical cumulate textures (Figure 4.2.1.9; Blackman et al., 2006). However, several lines of evidence suggest that these rocks are not simply the first crystallized product of a closed, fractionating magma body. Trace element patterns calculated for melts in equilibrium with measured in situ plagioclase, clinopyroxene, and olivine compositions (Figure 4.2.1.9(A)) illustrate that the olivine is in disequilibrium with the MORB melts that crystallized the plagioclase and clinopyroxene (Drouin et al., 2009). Mineral chemistry, textural relations, and crystallographic preferred orientations indicate that the olivine-rich troctolites are the end product of melt–rock interaction between an initial olivine-bearing rock and MORB melt (Drouin, Godard, Ildefonse, Bruguier, & Garrido, 2009; Drouin, Ildefonse, & Godard, 2010). The observed, relatively stronger concentrations of olivine [001] axes in some of these rocks are interpreted as resulting from melt– rock interaction. A first stage of dunitization weakens the olivine crystallographic preferred orientation (e.g., Tommasi, Godard, Coromina, Dautria, & Barsczus, 2004). It is followed by the disaggregation of mantle peridotite
of Hole 735B; the “strip average” estimate was obtained using a strip-sampling technique (see Hart et al., 1999) to take into account alteration, small-scale lithologic variability, and patchiness of veins, and thus obtain a possibly more representative sampling. The gray field indicates values of primary MORB melts which represent the most primitive basaltic compositions expected at a spreading centre (data from Kinzler and Grove (1993) in (A) and from Workman and Hart (2005) in (B)). Back stripping the contribution of olivine at Hole U1309 gives a composition for the bulk gabbroic crust at Hole U1309 consistent with estimates of the Mg#–Ca# composition of primary MORB magmas (BG). However, this estimate is characterized by trace element signatures indicative of plagioclase accumulation. Only models taking into account the contribution of leucocratic dykes in the bulk hole estimates give results consistent with primary magma compositions for both major and trace elements.
476 Earth and Life Processes Discovered from Subseafloor Environments
(A)
?
(B)
(C)
FIGURE 4.2.1.9 Sample characteristics from olivine-rich troctolite intervals (modified from Blackman et al., 2011). (A) Computed rare earth element compositions of melts in equilibrium with the measured mineral chemistry (Drouin et al., 2009). (B) Rounded and subhedral olivine crystals (Ol) included in large clinopyroxene (Cpx) and plagioclase poikiloblasts forming poikilitic texture in olivine-rich troctolite (sample 345_U1309D_247R3_16–18). (C) Coarse‐grained subhedral olivine crystals with smooth edges and well-defined and widely spaced subgrain boundaries and poikiloblastic plagioclase (Pl) and orthopyroxene (Opx) (sample 345_U1309D_248R2_22–24). (B) and (C) are cross-polarized photomicrographs from Drouin et al. (2010).
(Figure 4.2.1.10), which, if not heavily fluxed with melt, would be expected to display [100] preferred alignment related to crystal–plastic deformation (Drouin et al., 2010). Disruption of preexisting high-temperature crystallographic preferred orientations during melt influx is also supported by the common occurrence of adjacent grains with similar crystallographic orientations (Drouin et al., 2010; Suhr et al., 2008). Although not unequivocally demonstrated, a mantle origin for the olivine-rich troctolites is favored (Figure 4.2.1.10; Drouin et al., 2009, 2010; Suhr et al., 2008). This proposed scenario suggests a new mechanism for the formation of the lowermost oceanic crust in which small volumes (approximately 1–10-m-thick intervals in the core) of mantle peridotite are impregnated by, and react with, basaltic melts until they become isolated and incorporated into the crust during protracted periods of construction by igneous intrusion. Whatever the initial lithology, impregnation by large volumes of melt has strongly modified the original composition and microstructure of the host rock, and a melt/rock ratio of 3:1 is estimated, based on a fractionation model to match the Mg, Ni, and Cr contents measured in olivine grains (Suhr et al., 2008). The few recovered intervals of mantle peridotites (see above) also show petrographic and geochemical evidence of melt–rock reactions (Tamura et al., 2008; Godard et al., 2009; von der Handt & Hellebrand, 2010).
Formation and Evolution of Oceanic Lithosphere Chapter | 4.2.1 477
MORBtype melt
Ol-rich troctolites [100]
ridge axis
lithosphere
[010]
[001]
Ol
Impregnation : - Melt fraction ≥ ~30% - Disruption of olivine framework & modification of CPO
2
Dunite Ol
[100]
[010]
[001]
asthenosphere
Dunitization : Weakening of original CPO
1
[100]
[010]
[001]
Ol
Opx
Harzburgite FIGURE 4.2.1.10 Sketch of the proposed scenario for the formation of Hole U1309D olivine-rich troctolites (Drouin et al., 2010). Percolation of basaltic melts through depleted mantle peridotite at near-solidus temperature results in dunitization (orthopyroxene dissolution, precipitation of olivine, and formation of dunite), with weakening of olivine crystallographic preferred orientation (CPO) (indicated by 1), and impregnation when increasing melt fraction (≥30%) locally disrupts the olivine framework and further modifies the olivine CPO (indicated by 2). The dunitization and impregnation are viewed as successive stages of a continuous process. As temperature decreases, melt crystallizes plagioclase and clinopyroxene within the assimilated olivine matrix to form o livine‐rich troctolite. The schematic cross-section of the ridge axis is modified from Cannat (1996). The examples of original harzburgite and dunite CPOs are from Polynesian xenoliths (RPA18A and UAH280X, respectively; Tommasi et al., 2004). Note the absence of structural reference frame (no visible foliation in samples) for the dunite and olivine-rich troctolite CPOs.
4.2.1.3 DEEP DRILLING OF INTACT OCEAN CRUST FORMED AT A SUPERFAST SPREADING RATE: HOLE 1256D 4.2.1.3.1 Background The vast majority (∼70%) of magma derived from the mantle is brought into the Earth’s crust at mid-ocean ridges, and approximately two-thirds of that magma cools and crystallizes in the lower portion of the oceanic crust. Although <20% of modern ridges are spreading at fast rates, nearly 50% of present-day ocean crust (i.e., ∼30% of the Earth’s surface) was formed at fast-spreading ridges. Because fast-spread crust is more continuous and regularly layered than slowspread crust, extrapolating fast-spreading accretion processes from a few sites
478 Earth and Life Processes Discovered from Subseafloor Environments
might reasonably describe a significant portion of the Earth’s surface. Additionally, the great majority of crust subducting into the mantle over the past ∼200 my formed at fast-spreading ridges (e.g., Müller et al., 2008), making characterizing this style of crust most relevant for understanding the recycling of crustaland ocean-derived components back into the mantle. Drilling a complete in situ section of ocean crust has been an unfulfilled ambition of Earth scientists for many decades, which provided the impetus for the conception of scientific deep ocean drilling (e.g., Teagle & Ildefonse, 2011); ODP Hole 1256D is the latest, and to date most successful attempt to reach this goal. The deep drilling campaign to Site 1256 was designed to understand the formation, architecture, and evolution of ocean crust formed at fast spreading rates, and has been the focus of four scientific ocean drilling cruises (ODP 206, Wilson et al., 2003; IODP Expedition 309/312, Teagle et al., 2006; IODP Expedition 335, Teagle et al., 2012). Hole 1256D is located in 3635 m of water in the Guatemala Basin (6°44.2′N, 91°56.1′W), on the Cocos plate in the eastern equatorial Pacific Ocean. Ocean crust here formed 15 my ago during a sustained episode of superfast ocean ridge spreading (>200 mm per year; Wilson, 1996; Figures 4.2.1.1 and 4.2.1.11) at the EPR. The site formed on a ridge segment is at least 400 km in length, ∼100 km north of the ridge–ridge–ridge triple junction between the Cocos, Pacific, and Nazca plates. Ocean crust formed at a superfast 500 km
Seafloor age (my) 0
5
10
15
20
20°N
472
100°W
159
erton
Clipp
10°N
495 80°W
FZ
845 844
Site 1256 852/1243 851
79
Hess Deep
83/503
Cocos plate
82
81/572 80
504 Pacific plate
Nazca plate
FIGURE 4.2.1.11 Age map of the Cocos plate and corresponding regions of the Pacific and Nazca plates (Wilson et al., 2003). Isochrons at 5 my intervals have been converted from magnetic anomaly identifications according to the timescale of Cande and Kent (1995). Selected DSDP and ODP sites that reached basement are indicated by circles. The wide spacing of the 10–20 my isochrons to the south reflects the extremely fast (200–220 mm per year) full spreading rate.
Formation and Evolution of Oceanic Lithosphere Chapter | 4.2.1 479
spreading rate was deliberately targeted because there is strong evidence from mid-ocean ridge seismic experiments that gabbros occur at shallower depths in intact ocean crust with higher spreading rates (Carbotte, Mutter, Mutter, & Ponce-Correa, 1998; Purdy, Kong, Christeson, & Solomon, 1992; Wilson et al., 2006). Consequently, the often difficult-to-drill upper ocean crust should be relatively thin. Site 1256 has a seismic structure reminiscent of typical Pacific offaxis seafloor (Figure 4.2.1.12(B)). Upper Layer 2 velocities are 4.5–5 km/s, and the Layer 2–3 transition is between ∼1200 and 1500 m subbasement. The total crustal thickness at Site 1256 is estimated at approximately 5–5.5 km. Site 1256 sits atop a region of smooth seafloor and basement topography (<10 m relief; Figure 4.2.1.12). Unpublished processing (A.J. Harding, personal communication; Teagle et al., 2006) reveals discernible variation in the average seismic velocity (approximately 4.54–4.88 km/s) of the uppermost (∼100 m) basement. Two principal features are apparent (Figure 4.2.1.12(A)): a 5–10-km-wide zone of relatively high upper basement velocities (>4.82 km/s) that can be traced ∼20 km to the edge of data coverage southeast of Site 1256, and a bull’s eye of relatively low velocity (4.66–4.54 km/s) centered around the crossing point of two seismic Lines (21 and 25; Figure 4.2.1.12). The only serious drawback to the Site 1256 area as a crustal reference section for fast spreading rates is its low paleolatitude, making the determination of magnetic polarity from azimuthally unoriented core samples very challenging. In addition, the nearly north–south ridge orientation at the time of accretion makes the magnetic inclination relatively insensitive to tilting.
4.2.1.3.2 Summary of Upper Crustal Lithology at Site 1256 Hole 1256D was the first scientific borehole prepared for deep drilling in ocean crust by installing a large reentry cone secured with almost 270 m of 16-inch casing through the 250-m-thick sedimentary overburden and cemented into the uppermost basement. It was then deepened through an ∼810-m-thick sequence of lavas and a thin (∼346 m) sheeted dike complex, the lower 60 m of which is strongly contact metamorphosed to granoblastic textures. The first gabbroic rocks were encountered at 1407 mbsf (IODP Expedition 312), where the hole entered a complex dike–gabbro transition zone that includes two 20–50-m-thick gabbro lenses intruded into dikes with granoblastic textures (Figures 4.2.1.13 and 4.2.1.14). At the end of IODP Expedition 312, Hole 1256D had a total depth of 1507 mbsf and was open to its full depth. Five and a half years later, IODP Expedition 335 aimed to deepen Hole 1256D several hundred meters into the cumulate gabbroic rocks of intact lower oceanic crust to address fundamental scientific questions about lower crustal accretion and cooling mechanisms. Unfortunately Hole 1256D was deepened only minimally (to 1521 mbsf), as a number of significant engineering challenges were encountered and overcome during the cruise, each unique but not unexpected in a deep, uncased marine borehole into igneous rocks (see Operations section of the Expedition Summary in Teagle et al., 2012).
(B)
FIGURE 4.2.1.12 Characteristics of the Site 1256 area (modified from Teagle et al., 2012). (A) Bathymetry; abyssal hill relief of as much as 100 m is apparent in the southwest part of the area; relief to the northeast is lower and less organized. Numbers 21–28 identify multichannel seismic (MCS) lines (white) from the site survey. Red (light gray in print versions) and blue (dark gray in print versions) contours are selected top-of-basement seismic velocity contours; the red contour encloses velocity >4.82 km/s, which is interpret as a plausible proxy for the presence of thick ponded lava flows, as sampled in Hole 1256D (Teagle et al., 2012; Wilson et al., 2006); the blue contour encloses velocities <4.60 km/s, possibly reflecting a greater portion of pillow lavas than elsewhere in the region. (B) Site survey MCS profile (Hallenborg et al., 2003) from line 22 that crosses at Site 1256, with penetration as of the end of Expedition 335 scaled approximately to traveltime. Horizontal reflectors in the upper basement to traveltimes of 5.5 s appear to result from contrasts between lava flow sequences, corresponding to depths of at least 800 m in basement.
480 Earth and Life Processes Discovered from Subseafloor Environments
(A)
Formation and Evolution of Oceanic Lithosphere Chapter | 4.2.1 481
(A)
Lith. Sedimentary overburden
0
Recovery
Massive basalt Sheet flow Pillow basalt Sheeted dike Granoblastic dike Gabbro
40
250
Resistivity ( m)
1
10 102 103 104 105
(C)
0
Porosity (%) 5
10
15 4
(D)
Vp (km/s) 5
6
7
(E)
Inflated flows
Sheet and massive flows
Leg 206
70
Refraction velocity measurement Inversion of traveltimes from Leg 206 VSP Inversion of traveltimes from Expedition 312 VSP
Exp 312
1250
Sheeted dike complex
Exp 309
Depth (mbsf)
750
1000
60
(B)
Lava pond
500
Mg #
50
Intrusive contacts Magmatic contacts Breccia
1500 Junk Baskets
FIGURE 4.2.1.13 Downhole plots of data for ODP Hole 1256D (modified from Wilson et al., 2006). (A) Summary lithostratigraphic column of the basement drilled at Site 1256 showing recovery, and major lithologies. (B) Mg# obtained from shipboard and published whole rock measurements (Neo, Yamazaki, Miyashita, & the Expedition 309/312 Scientists, 2009; Teagle et al., 2006, 2012; Wilson et al., 2003; Yamasaki et al., 2009). (C) Electrical resistivity logging (deep) data (Teagle et al., 2012). (D) Porosity (Gilbert & Salisbury, 2011). Small black symbols are shipboard measurements on discrete samples, red (dark gray in print versions) symbols are postcruise laboratory measurements, and the light blue (pale gray in print versions) curve is calculated from the DLL (Dual Laterolog tool) deep electrical resistivity log. (E) Compressional seismic velocity (modified from Gilbert & Salisbury, 2011). Small black symbols are shipboard measurements on discrete samples corrected for lithostatic pressure, red (dark gray in print versions) symbols are postcruise laboratory measurements under lithostatic pressure, the blue (light gray in print versions) curve is the sonic log data, as reprocessed by Guérin, Goldberg, and Iturrino (2008), the dashed red (dark gray in print versions) line is derived from seismic refraction data (Wilson et al., 2003), the bold red (dark gray in print versions) and orange (light gray in print versions) lines show the inversion of traveltimes from Leg 206 and Expedition 312 vertical seismic profiles (VSPs), respectively (Swift et al., 2008).
However, IODP Expedition 335, returned a unique collection of samples from the granoblastic dikes (Figures 4.2.1.13 and 4.2.1.14), obtained from the many fishing, reaming, and cleaning runs with junk baskets. Some of these samples were much larger than normal core pieces, up to ∼20 cm in equivalent diameter.
(A)
Olivine gabbro
Gabbro
Gabbronorite
100%
Oxide gabbro
1400
0%
Diorite/Tonalite
Recovery
Granoblastic dike
482 Earth and Life Processes Discovered from Subseafloor Environments
(B)
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FIGURE 4.2.1.14 Plutonic section from the lower portion of Hole 1256D with representative photomicrographs of key samples (modified from Wilson et al., 2006; Teagle et al., 2012). The distribution of rock types is expanded proportionately in zones of incomplete recovery. (A) Photomicrograph of a dike completely recrystallized to a granoblastic association of equant secondary plagioclase, clinopyroxene, magnetite, and ilmenite (plane-polarized light). (B) Uppermost dike/gabbro boundary; medium-grained oxide gabbro is intruded into a granoblastically recrystallized dike along an irregular, moderately dipping contact. (C) Quartz-rich oxide diorite strongly altered to actinolitic hornblende, secondary plagioclase, epidote, and chlorite. (D) Disseminated oxide gabbro with patchy texture and centimeter-scale dark ophitically intergrown clinopyroxene and plagioclase patches separated by irregular, more highly altered leucocratic zones. (E) Sharp modal contact between a medium-grained olivine-rich gabbro (top in the core) and a gabbro. (F) Clast of partially resorbed granoblastic dike within gabbro. (G) Photomicrograph of a granoblastic basalt (plane-polarized light) showing a relict plagioclase microphenocryst with pyroxenes and oxide inclusions in its core (1) and a metamorphic, orthopyroxene vein (2). (H) Dioritic vein and
Formation and Evolution of Oceanic Lithosphere Chapter | 4.2.1 483
The volcanic sequences at Site 1256 consist of massive flows and thinner sheet flows with subordinate pillow basalt, hyaloclastite, and breccia ( Figure 4.2.1.13). The uppermost crust at Site 1256 comprises an ∼100-m-thick sequence of lava dominated by a single massive lava flow >75-m thick, requiring at least this much seafloor relief to pond the lava (Tartarotti & Crispini, 2006; Wilson et al., 2003). On modern fast-spreading ridges, such topography does not commonly develop until 5–10 km from the axis (Macdonald, Fox, Alexander, Pockalny, & Gente, 1996). Although this lava flow cooled off-axis it may have originated at the ridge axis before flowing on to the ridge flanks, as is observed for very large lava flows on modern ocean floor (e.g., Macdonald, Haymon, & Shor, 1989). This flow is relatively unfractured at the scale of the drill hole (Crispini, Tartarotti, & Umino, 2006), with shipboard physical property measurements on discrete samples indicating VP > 5.5 km/s (Wilson et al., 2003). Hence, the area of relatively high uppermost basement seismic velocities likely represents the regional extent of this massive flow, although this hypothesis has not been tested by drilling. Assuming an average thickness of 40 m, the regional survey results would conservatively suggest an eruption volume in excess of 3 km3, plausibly >10 km3. Lava flows this voluminous are not described on-axis at fast-spreading ridges. An off-axis origin is suggested by the presence of a strongly welded and plastically deformed agglutinate in the base of the thick lava in Hole 1256C, that indicates proximity to the source vent for the lava field (Crispini et al., 2006). An off-axis volcanic flow field of comparable size is described on the EPR at 14°S (Geshi et al., 2007). This is extremely large when compared to the size of mid-ocean ridge axial lowvelocity zones that are thought to be high-level melt lenses, which typically have volumes approximately 0.05–0.15 km3 per kilometer of ridge axis, and generally appear to be only partially molten (Singh, Kent, Collier, Harding, & Orcutt, 1998). This poses questions regarding the melt sources of these very large submarine flow fields that require either the draining of a relatively large magma reservoir so far not imaged at fast-spreading mid-ocean ridges (Teagle et al., 2012), or a magma system that is being recharged during the course of slow (several decades long), ongoing eruptions, such as described in Iceland (e.g., Eason & Sinton, 2009). The lava sequence below the lava pond includes sheet and massive flows and minor pillow flows. Subvertical, elongate flow-top fractures filled with quenched glass and hyaloclastite in these lavas indicate flow lobe inflation requiring cooling on a subhorizontal surface off-axis (Umino, Lipman, & Obata, 2000). It is estimated that a total thickness of 284 m of lava flowed and cooled off-axis. Sheet
patch within granoblastic basalt. (I) Sharp, curviplanar dike/dike contact between a light-colored, fine-grained basalt (right) and a darker, coarse-grained basalt (left). (J) Medium-grained, orthopyroxene-bearing olivine gabbro.
484 Earth and Life Processes Discovered from Subseafloor Environments
flows and massive lava flows erupted at the ridge axis make up the remaining extrusive section to 1004 mbsf. This contrasts slightly with the volcanic stratigraphy for Hole 1256D developed from analysis of wireline geophysical imaging (Tominaga & Umino, 2010; Tominaga, Teagle, Alt, & Umino, 2009) that suggests <50% of the lava drilled crystallized within 1000 m of the axis but that the majority of the lava pile had formed within 3000 m of the ridge crest. From 1004 to 1061 mbsf is the Transition Zone between the extrusive rocks and the sheeted dike complex, marked by the occurrence of subvertical intrusive contacts, a silica-sulfide mineralized hyaloclastic breccia zone, and a step change in hydrothermal alteration from low temperature to greenschist facies assemblages. Below this transition, numerous subvertical intrusive contacts were observed, indicating the start of a relatively thin, ∼350-m-thick sheeted dike complex dominated by massive basalt. Some basalts have doleritic textures, and many are crosscut by subvertical dikes with strongly brecciated chilled margins that are mineralized by hydrothermal fluids. There is little evidence from core or geophysical wireline logs for major tilting of the dikes, consistent with subhorizontal seismic reflectors in the lower extrusive rocks that are continuous for several kilometers across the site (Hallenborg, Harding, Kent, & Wilson, 2003). Measurements of dike orientations using FMS images are in close agreement with direct measurements on recovered cores and indicate that the dikes at Site 1256 are tilted ∼15° away from the paleospreading axis around a ridgeparallel, subhorizontal axis (Tominaga et al., 2009; Violay, Pezard, Ildefonse, Célérier, & Deleau, 2012). In the bottommost ∼60 m of the sheeted dikes (1348–1407 mbsf), basalts are partially to completely recrystallized to distinctive granoblastic textures characterized by granular secondary clinopyroxene and lesser orthopyroxene, resulting from contact metamorphism by underlying gabbroic intrusions (Figures 4.2.1.13 and 4.2.1.14). Textural changes in oxide minerals indicate that the zone of metamorphic recrystallization may extend for >90 m above the upper contact with the first gabbro interval (Coggon, Alt, & Teagle, 2008). There is unequivocal evidence for hydrothermal alteration before and after the formation of the granoblastic textures and potentially small, irregular patches of partial melting (Alt et al., 2010; France, Ildefonse, & Koepke, 2009; Koepke et al., 2008; Teagle et al., 2006). Simple thermal calculations suggest that the two gabbro bodies so far intersected in Hole 1256D have insufficient thermal mass to be responsible for a 60–90-m-thick high-temperature (600–900 °C) contact-metamorphic halo if the intrusions are simple subhorizontal bodies (Coggon et al., 2008; Koepke et al., 2008). The amount of heat needed to produce this metamorphism requires either a much larger intrusion nearby, or significant topography on the dike/gabbro boundary (Teagle et al., 2006). Gabbro and trondhjemite intrusions into sheeted dikes at 1407 mbsf mark the top of the plutonic section (Figures 4.2.1.13 and 4.2.1.14) and the start of a complex dike–gabbro transition zone. This first occurrence of plutonic rocks from intact ocean crust occurred within the depth range predicted from
Formation and Evolution of Oceanic Lithosphere Chapter | 4.2.1 485
the relationship between spreading rate and depth to axial low-velocity zones (Teagle et al., 2006). Two major bodies of gabbro were penetrated beneath this contact, with the 52-m-thick upper gabbro (Gabbro 1) separated from the 24-m-thick lower gabbro (Gabbro 2) by a 24-m screen of granoblastic basalts (Dike Screen 1). The textures and rock types observed in Hole 1256D gabbro intervals are similar to those of varitextured gabbros thought to represent a frozen melt lens between the sheeted dike complex and the underlying cumulate gabbros in ophiolites (e.g., MacLeod and Yaouancq, 2000; France et al., 2009). Gabbro 1 is mineralogically and texturally heterogeneous, comprising olivine gabbros, gabbros, oxide gabbros, quartz-rich oxide diorites, and small trondhjemite dikelets. Oxide abundance decreases irregularly downhole, and olivine is present in significant amounts in the lower part of Gabbro 1. The intervening Dike Screen 1 is interpreted as an interval of sheeted dikes captured between the two intrusions of gabbros. It consists of fine-grained meta-basalt similar to the granoblastic dikes overlying Gabbro 1, and is cut by a number of small quartz gabbro and tonalite dikelets of variable thickness (1–10 cm), grain size, and composition. Gabbro 2 comprises gabbro, oxide gabbro, and subordinate orthopyroxene-bearing gabbro and trondhjemite that are altered similarly to Gabbro 1, and has clear intrusive contacts with the overlying granoblastic Dike Screen 1. Partially resorbed, stoped dike clasts are entrained within both the upper and lower margins of Gabbro 2 (Figure 4.2.1.14). Gabbro 2 is characterized by the absence of fresh olivine, high but variable orthopyroxene contents (5–25%), and considerable centimeter- to decimeter-scale heterogeneity. Oxide abundance generally diminishes downhole. The predominant rock type is orthopyroxene-bearing gabbro with gabbronorite in the marginal units. Dike Screen 2, at the bottom of the hole, was extensively sampled with junk baskets during IODP Expedition 335 (Teagle et al., 2012). It is dominantly made of fine-grained granoblastic aphyric basalts, most of which are strongly recrystallized (Figure 4.2.1.14). Some samples preserve dike-to-dike contacts, and many preserve cores of former plagioclase (micro)phenocrysts (Figure 4.2.1.14), indicating that the granoblastic rocks are contact-metamorphosed sheeted dike diabase. Approximately half of the granoblastic basalt samples contain small, irregular patches (<5 × 3 cm), veins (approximately 1–2 mm wide), and dikelets (<1.5 cm wide) of evolved plutonic rocks (oxide gabbro, diorite, and tonalite) (Figures 4.2.1.14 and 4.2.1.15). The veins are observed to be offshoots of the igneous patches, and diffuse patches are issued from the dikelets. Hence these features form part of the same network, marking a single generation of intrusion of melts into the granoblastic basalts. Their magmatic textures (subhedral to euhedral shapes of several phases, along with poikilitic textures) contrast strongly with the granoblastic textures of the host rocks, demonstrating that intrusion occurred after the high-temperature recrystallization of the host rock. The common occurrence of primary magmatic amphibole in the veins and patches suggests high water activities during their formation, hence possibly the partial
486 Earth and Life Processes Discovered from Subseafloor Environments
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FIGURE 4.2.1.15 Schematic “outcrop” of the dike–gabbro transition zone, showing the micro-textures and contact relationships observed in the cores and junk basket samples recovered during Expedition 335 (Teagle et al., 2012). (A) Diorite dikelet, comprised predominantly of primary igneous amphibole and plagioclase, crosscutting granoblastic basalt (Sample 335-1256D-Run19-RCJB-Rock C). (B) Fine-grained dike chilled against coarse-grained dike (Sample 335-1256D-Run12-RCJB-Rock S). The entire sample is recrystallized to granoblastic assemblages of plagioclase, orthopyroxene, clinopyroxene, magnetite, and ilmenite. Later fine postcontact-metamorphic hydrothermal amphibole veins (not shown) cut across the contact and both dikes. (C) Chilled and brecciated dike margin recrystallized to granoblastic assemblages (Sample 335-1256D-Run14-EXJB-Foliated). Angular clasts consist of granoblastic plagioclase with minor orthopyroxene, clinopyroxene, ilmenite, and magnetite and are recrystallized from chilled dike margin breccia protolith. Clast matrix is orthopyroxene-rich with minor clinopyroxene, plagioclase, and oxides; it is recrystallized from hydrothermal minerals (amphibole and chlorite) that veined and cemented the breccia protolith. (D) Subophitic texture in gabbro (Sample 335-1256D-Run11EXJB). (E) Diorite vein crosscutting a conjugate set of metamorphic orthopyroxene veins (Sample 335-1256D-Run19-RCJB-Rock B). (F) Postcontact-metamorphic hydrothermal hornblende vein cutting granoblastic basalt (Sample 335-1256D-238R-1, 2–4 cm). (G) Granoblastic basalt with a diorite patch (Sample 335-1256D-Run11-EXJB-Rock). (H) Granoblastic orthopyroxene vein, recrystallized from precursor hydrothermal vein, in granoblastic dike (Sample 335-1256D-238R-1, 2–4 cm). Orthopyroxene vein is cut by small postcontact-metamorphic hydrothermal amphibole vein. (I) Postcontact-metamorphic hydrothermal amphibole vein cutting granoblastic basalt, with 1-mm-wide amphibole-rich alteration halo where pyroxenes are replaced by amphibole (Sample 335-1256D-Run12-RCJB-Rock B).
melting of previously hydrothermally altered rocks. Moreover, the presence of quartz, as well as accessory apatite and zircon, implies that patches, veins, and dikelets crystallized from highly evolved melts. Only a small number of gabbroic rocks (∼3 wt%) were recovered during IODP Expedition 335
Formation and Evolution of Oceanic Lithosphere Chapter | 4.2.1 487
(Figure 4.2.1.14) and these rocks range in composition from disseminated oxide gabbro to orthopyroxene-bearing olivine gabbro. The IODP Expedition 335 gabbroic rocks are texturally less variable than those recovered during IODP Expedition 312 from the gabbro 1 and 2 intervals. Although the contact relationships with the granoblastic basalts were not recovered, it is likely that at least some of the gabbroic rocks occur intercalated with the granoblastic basalt, perhaps forming small intrusions. The unique sampling with fishing tools achieved during IODP Expedition 335 allowed the documentation of a variety of intrusive structural/textural relationships, and overprinting/crosscutting hydrothermal alteration and metamorphic paragenetic sequences that had not previously been observed in situ, owing to the one-dimensional nature of drill cores and the very low rates of recovery of the granoblastic material in the sheeted dike/gabbro transition zone. Overall, the bottom of Hole 1256D documents a section of metamorphosed, granoblastic sheeted dikes that underwent small-scale intrusion by both gabbroic and evolved plutonic rocks (Figure 4.2.1.15).
4.2.1.3.3 A Thin upper Crust at Superfast Spreading Rate Relative to other well-studied upper ocean crust sections (e.g., Karson, 2002), Site 1256 is unique as it shows a thick lava sequence and a thin dike sequence (Umino et al., 2008). Steady-state thermal models require that the conductive lid separating magma from rapidly circulating seawater gets thinner as spreading rate increases, indicating that the thin dike sequence is a direct consequence of the high spreading rate. A thick flow sequence, with many massive individual flows and few pillow lavas, is a consequence of the elevated magma budget (Umino et al., 2008) and/or of the short vertical transport distance from the magma chamber, and is similar to observations from the midsegments of the fast-spreading southern EPR (White, Macdonald, & Haymon, 2000). The high lava/dike thickness ratio appears to be characteristic of superfast spreading, which is in direct contrast to spreading models developed from observations of tectonically disrupted fast-spread crust exposed in Hess Deep (Karson, 2002) that suggest regions of high magma supply should have thin lavas and thick dikes. Similarly, there is little evidence for tilting (at most ∼15°) in Hole 1256D and no evidence for significant faulting. In contrast, the upper crust exposed at Hess Deep (fast spreading) and in the Blanco Fracture Zone (intermediate spreading) shows significant faulting and rotation within the dike complex (Karson, 2002), indicating that observations from those tectonic windows may be site specific and not widely applicable to intact ocean crust. The massive lava flow at the top of the Site 1256 basement indicates that faults with throws of approximately 50–100 m must exist in superfast spreading rate crust relatively near to the ridge axis (<10 km) in order to provide the relief necessary to pond the lava (as described at the southern EPR; e.g., Auzende et al., 1996; Macdonald, 1998; Sinton et al., 2002).
488 Earth and Life Processes Discovered from Subseafloor Environments
4.2.1.3.4 Where Is the Layer 2/3 Transition at Site 1256? Marine seismologists traditionally subdivided the ocean crust into seismic layers: Layer 1 comprises low-velocity sediments (VP < 3 km/s); Layer 2 has low velocity and a high velocity gradient, with VP typically ranging from ∼3.5 to ∼6.7 km/s; and Layer 3 has high velocity and a low velocity gradient (VP ranges from 6.7 to 7.1 km/s); the step-up to VP > 8 km/s marking the Mohorovičić Discontinuity and is interpreted to be the boundary with ultramafic rocks of the uppermost mantle. In casual language “Layer 3” is often mistakenly used as a synonym for gabbro, although previous drilling in Hole 504B has penetrated Layer 3 but not yet reached gabbroic rocks (Alt et al., 1996; Detrick et al., 1994). From regional seismic refraction data, the transition from seismic Layer 2 to Layer 3 at Site 1256 occurs between 1200 and 1500 m into basement (Wilson et al., 2003). Shipboard determinations of seismic velocities of discrete samples, once corrected for lithostatic pressure, are in close agreement with postcruise laboratory measurements of discrete samples under lithostatic pressure, and with in situ measurements by wireline tools (Figure 4.2.1.13; Gilbert & Salisbury, 2011). Contrary to expectation, porosity increases and P-wave velocities decrease stepwise downward from the lowermost dikes into the uppermost gabbro in Hole 1256D, as the result of the contact metamorphism of the granoblastic dikes and the strong hydrothermal alteration of the uppermost gabbros (Figure 4.2.1.13). Porosity and velocity then increase downhole in the gabbro. Wireline velocity measurements end at the top of gabbro, because additional velocity data closer to the bottom of the hole could not be acquired at the end of IODP Expedition 335 (Teagle et al., 2012). The examination of shipboard and postcruise discrete sample measurements, wireline logging, and VSP data suggests that the base of Hole 1256D has just entered, or is close to the transition from Layer 2 to Layer 3 (Gilbert & Salisbury, 2011; Swift, Reichow, Tikku, Tominaga, & Gilbert, 2008). Encountering gabbro at a depth within Layer 2 reinforces previous inferences that factors such as porosity and hydrothermal alteration (Alt et al., 1996; Carlson, 2010; Detrick et al., 1994) are more important than rock type or grain size in controlling the location of the Layer 2–3 transition. Drilling deeper in Hole 1256D is required to characterize the Layer 2–3 transition at Site 1256.
4.2.1.3.5 Geochemical Characteristics of the Upper Crust at Site 1256 Flows and dikes from Hole 1256D are N-MORB but show a wide range of magmatic fractionation (Figures 4.2.1.13 and 4.2.1.16). The lavas and dikes are on average relatively fractionated compared to global MORB data compilations as anticipated for fast-spreading EPR MORB (e.g., Rubin & Sinton, 2007). The lava pond, cored in both Holes 1256C and D, is more fractionated (Mg# 50 ± 2) than the rocks from the inflated flows, sheet and massive flows, and the sheeted
Formation and Evolution of Oceanic Lithosphere Chapter | 4.2.1 489 6KHHWIORZ
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0J2 ZW FIGURE 4.2.1.16 FeOT (Total Fe expressed as FeO) versus MgO for the basement at Site 1256 (Shipboard data from Leg 206 and Expeditions 309, 312, and 335; Teagle et al., 2006, 2012). Data are compared with analyses of northern East Pacific Rise (purple contour; Langmuir, Bender, & Batiza, 1986). Dashed lines show constant Mg#. Possible primary mantle melt compositions should have Mg# of 70–78 and MgO of 9–14 wt%. All flows and dikes, and most gabbros are too evolved to be candidates for primary magmas.
dike complex that all show a similar range of compositions (Mg# 56 ± 6). The more differentiated nature of the lava pond lavas is also displayed by relatively high incompatible element (e.g., Zr, TiO2, Y, and V) and low compatible element (Cr and Ni) concentrations (Teagle et al., 2006; Wilson et al., 2003). An interval within the lava pond cored in Hole 1256C (294–306 mbsf) shows anomalously high K2O concentrations (∼0.74 wt%) that are difficult to explain by magmatic fractionation and are not a result of hydrothermal alteration by seawater. There are variations in the basalt chemistry downhole, with a step in igneous trends at 750 mbsf possibly indicating a cycle of fractionation, replenishment, and, perhaps, assimilation (Figure 4.2.1.13). Downhole geochemical compositions within the dike section are variable and do not define trends; primary and evolved compositions are closely juxtaposed, as would be expected for vertically intruded magmas (Figure 4.2.1.13). The range of major element compositions in the dikes is similar to that of the overlying rocks, and the average composition of the dikes is indistinguishable from the average composition of the lavas (Figures 4.2.1.13 and 4.2.1.16). The compositional ranges of fresh lava and dike samples correspond to typical values for N-MORB for most major elements and many trace elements (Gale, Dalton, Langmuir, Su, & Schilling, 2013; Rubin & Sinton, 2007) and are
490 Earth and Life Processes Discovered from Subseafloor Environments
similar to those observed for the northern EPR (Figure 4.2.1.16). When trace elements are compared to EPR MORB, they are within one standard deviation of average, albeit on the relatively trace element-depleted side of MORB. For example, compared with first-order mid-ocean-ridge segments along the EPR, basalts from Site 1256 have low Zr/TiO2 and Zr/Y. This decrease in trace element ratios would require ∼30% more melting at the superfast ridge, but this appears unlikely because of normal crustal thickness (∼5.5 km) at Site 1256. More likely, the relatively depleted trace element signature at ODP Site 1256 results from the upwelling of previously depleted Galápagos plume mantle at the paleo-Site 1256 ridge axis (Geldmacher et al., 2013; Park, MacLennan, Teagle, & Hauff, 2008). The gabbro compositions, as documented by shipboard geochemical analyses, span a range similar to the flows and dikes but have on average higher MgO and lower FeO, albeit still within the range of EPR basalts (Figure 4.2.1.16). The uppermost gabbros have geochemical characteristics similar to the overlying dikes with MORB chemistries (MgO approximately 7–8 wt%; Zr approximately 47–65 ppm). Deeper, in Gabbro 2, the rocks are less fractionated and there are general downhole trends of increasing MgO, CaO, and Ni and decreasing FeO, Zr, and Y. Decreasing concentrations of FeO and TiO2 downhole suggest that Gabbro 2 is fractionated similarly to Gabbro 1. The intrusive nature of both gabbro bodies and the chemical variations within them suggest that they intruded into the base of the sheeted dikes and underwent minor internal fractionation in situ, resulting in the observed general geochemical stratification. Partially resorbed xenoliths of granoblastic dikes within Gabbro 2 (Figure 4.2.1.14) indicate that stoping of the intruded dikes may have contaminated the gabbro compositions (e.g., France et al., 2009). There are linear trends in the gabbro units between MgO and TiO2, FeO, CaO, Na2O, and Zr, most likely resulting from fractional crystallization of a gabbroic magma, with the fractionating assemblage consisting of clinopyroxene, plagioclase, and subordinate olivine, as expected for relatively evolved basaltic magmas. Olivines within the gabbros have Forsterite composition Fo < 81 (Yamazaki, Neo, & Miyashita, 2009) indicating fractionation. Similarly, even the most primitive dikes and lavas at Site 1256 (Mg# ∼66) would be in equilibrium with olivine more evolved than mantle olivine, and hence must have undergone fractionation. Simple mass balance calculations indicate that the average basalt in Hole 1256D has lost >30% of its original liquid mass as solid gabbro, implying the presence of at least 300 m of cumulate gabbro in the crust below the present base of the hole (Teagle et al., 2006). Therefore, the residue removed from primary magma to produce the observed gabbro and basalt compositions must occur below the present base of Hole 1256D. The lower part of Hole 1256D, below the dike–Gabbro 1 boundary at ∼1406 mbsf, is characterized by large chemical variations with, for example, Mg# ranging from 42 to 72 (Figure 4.2.1.13), and Zr from 23 to 117 ppm (Teagle et al., 2012). This variability mainly reflects changes in rock types from the
Formation and Evolution of Oceanic Lithosphere Chapter | 4.2.1 491
low-Mg#, trace element-rich sheeted and granoblastic dikes and dike screens, to the higher Mg# and trace element-depleted gabbroic rocks of Gabbro 1 and Gabbro 2. There are significant depth-dependent trace element variations in the granoblastic basalts, with a general downhole trend of decreasing incompatible element contents (i.e., Zr and, to a lesser extent, Y). Details of these variations and their possible significance are presented in Teagle et al. (2012).
4.2.1.4 SHALLOW DRILLING IN FAST-SPREAD LOWER CRUST AT HESS DEEP The primary goal of the “superfast campaign” at Site 1256 was to reach gabbroic rocks of the lower crust. In spite of the operational difficulties encountered during IODP Expedition 335, Hole 1256D has been thoroughly cleaned from debris and cuttings, and is still open. Additional expeditions to Hole 1256D will be required to deepen the hole into the cumulate rocks of the lower crust. An alternative to deep drilling into intact, fast-spread oceanic crust (as at Site 1256) is to access the lower crust in the rare places where it crops out on the seafloor, such as in tectonic windows. This was the goal of the recent IODP Expedition 345, which targeted lower crust gabbros at Hess Deep in the equatorial eastern Pacific (Figure 4.2.1.1; Table 4.2.1.1; Gillis, Snow, Klaus, & the Expedition 345 Scientists, 2014a; Gillis et al., 2014b). The scientific objectives were to test endmember models of magmatic accretion and estimate the intensity of hydrothermal cooling at depth. The Hess Deep Rift was selected to exploit tectonic exposures of young EPR plutonic crust, building upon results from Ocean Drilling Program Leg 147 (Gillis et al., 1993) and submersible surveys (e.g., Francheteau et al., 1990; Hékinian, Bideau, Francheteau, Lonsdale, & Blum, 1993; Lissenberg, MacLeod, Howard, & Godard, 2013). Low to moderate recovery (15–30%) was achieved in three holes (35–110 mbsf). Olivine gabbro and troctolite represent 84% of the plutonic rock types recovered at Site U1415. Additional recovered lithologies are minor gabbro, clinopyroxene oikocryst-bearing troctolite, clinopyroxene oikocryst-bearing gabbro, and gabbronorite. All of these lithologies have Mg# between 78 and 90. Spectacular modal and/or grain-size layering, reminiscent of that classically described in ophiolites (e.g., Boudier, Nicolas, & Ildefonse, 1996; Kelemen, Koga, & Shimizu, 1997; Pallister & Hopson, 1981), was observed in more than half of the recovered material (Figure 4.2.1.17). Magmatic foliation is commonly moderate to strong, particularly in layered lithologies. Together with the compositions of gabbros from the upper part of the igneous crust (known from ODP Hole 894 samples; Gillis et al., 1993), and of the upper crust (known from numerous Alvin and Nautile dive samples; Coogan, Gillis, MacLeod, Thompson, & Hekinian, 2002, and references therein), Gillis et al. (2014b) calculated an estimate of the bulk composition of the Hess Deep crust, the first estimate for fast-spread crust, with an Mg# of 74. This average bulk composition can then be used as that of a parental melt for a MORB fractional crystallization model, which is consistent with the measured sample compositions (Gillis et al., 2014b). However, an unexpected finding, which
492 Earth and Life Processes Discovered from Subseafloor Environments
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PP
FIGURE 4.2.1.17 Examples of layered gabbro samples from IODP Hole U1415J at Hess Deep (Gillis et al., 2014a). (A) Core 345-U1415J-5R-2, piece 1. (B) 345-U1415J-8R-2, piece 9.
does not fit a simple fractional crystallization trend, is the common abundance (up to 5%) of orthopyroxene as an early cumulus phase in the rocks recovered during Expedition 345. This observation calls for a more complex petrogenetic history of the Hess Deep lower crust, which possibly involves reaction of ascending MORB melts with the upper mantle (hence delivery of melts to the lower crust that are not undersaturated in orthopyoxene; Coogan et al., 2002), and/or the contamination of parental melts by water (Boudier, Godard, & Armbruster, 2000; Nonnotte, Ceuleneer, & Benoit, 2005). The common occurrence of delicate skeletal olivine grains suggests that at this location, the lower crust was not significantly deformed after crystallization. IODP Expedition 345 provides a new, unique, reference section for fast-spreading lower crust (Gillis et al., 2014b), which complements the previous seafloor sampling in this area (Hékinian et al., 1993; Coogan et al., 2002; Lissenberg et al., 2013). Paleomagnetic demagnetization experiments conducted during IODP Expedition 345 on discrete samples of gabbros and troctolites yielded stable magnetization directions with a variety of remanence structures. In all samples, components that unblock close to the magnetite Curie temperature are considered to represent primary thermoremanent magnetizations acquired during crustal accretion. Downhole variations in the inclination of this stable, highunblocking temperature magnetic component indicate that at least two blocks have been sampled in each of the two deepest holes (U1415J and U1415P; Gillis, Snow,
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Klaus, & the Expedition 345 Scientists, 2014a). The data provide evidence of relative displacements of individual, internally coherent blocks (tens of meters in size), consistent with emplacement by very large-scale mass wasting on the southern slope of the Hess Deep intrarift ridge (Ferrini et al., 2013).
4.2.1.5 CONCLUSION During the last decade, IODP has made substantial progress in sampling of the ocean crust, by doubling the amount of deep (>1 km below seafloor) holes. Two deep holes from sections formed at extremely different spreading rates have shed light on very important unknown and/or poorly understood aspects of the architecture and accretion of modern ocean crust. In slow-spread crust, Hole U1309D was the second deep hole in an OCC following historic deep drilling at ODP Hole 735B (that remains the deepest hole in slow-spread crust, and returned the highest recovery in hard rock drilling). The spectacular igneous section recovered in the core of the Atlantis Massif is key to demonstrating the important role of magmatism in the development of OCCs. The interplay between magmatism and tectonics is still to be fully documented, described, and quantified, but Hole U1309D is a significant step forward. Olivine-rich troctolites, which represent about 5% of the material recovered in the hole, and which are by far the least altered olivine-rich material (locally <1% serpentinized) ever recovered by scientific ocean drilling, represent an exceptional record of the complex igneous accretion history, involving both fractional crystallization and melt–rock reactions, and possibly also assimilation of mantle peridotites. Another first, resulting from the combination of high recovery and high-quality borehole imaging (FMS) data, was the core-log integration that allowed reorientation of individual core pieces and associated paleomagnetic data and the provision of unequivocal constraints on the tectonic rotation associated with the development of the OCC. Hole 1256D, in Pacific superfast-spread crust, has for the first time reached the base of the sheeted dikes and penetrated the first gabbroic rocks ever sampled in intact, in situ layered fast-spread crust. The upper crustal section recovered at Site 1256 complements Hole 504B in intermediate-spread crust, which has been for a long time the single available reference section for the modern oceanic upper crust. The sampling of the root zone of the sheeted dike complex, and underlying varitextured gabbros and recrystallized granoblastic basalts, documents for the first time the transition zone between the present-day upper and lower crust formed at fast-spreading ridges. From this unique sampling emerges a vision of a complex, dynamic thermal boundary layer zone. This region of the crust between the principally hydrothermal domain of the upper crust and the intrusive magmatic domain of the lower crust is one of evolving geological conditions. The resulting intimate coupling among temporally and spatially intercalated magmatic, hydrothermal, partial melting, intrusive, metamorphic, and retrograde processes is recorded in the recovered samples.
494 Earth and Life Processes Discovered from Subseafloor Environments
In addition to the two deep holes in slow- and ultrafast-spread crust, another step forward was recently achieved by sampling the first significant sections of layered gabbros of “Penrose”-style ocean crust (Gillis et al., 2014b), assumed to constitute most of the lower crust based on ophiolite studies since the early 1970s. The drawback of accessing the deeper crustal levels in a tectonic window like Hess Deep is that the recovered sections are tectonically out of place, and hence cannot be used to fully document the depth variations of compositions and structures in the way required to test crustal accretion models (e.g., Teagle et al., 2012). Yet, this unique sampling of the lower, layered crust brings fundamental new constraints on the average composition of fastspread crust, on the petrological processes involved in building the layered gabbros, and in transporting melt from these deeper levels to the upper crust. Hence, critical samples and data recovered by scientific drilling over the past 10 years provide significant additional constraints on understanding lower oceanic crustal processes. However, we are still far from having a complete set of direct observations regarding the accretion occurring beneath the dike layer at fast-spreading ridges, and we need to achieve deep drilling of intact, fast-spread ocean crust. Deep drilling into ocean crust has posed, and will continue to present major technical and programmatic challenges to scientific ocean drilling. Only four holes, DSDP Hole 504B, ODP Holes 735B and 1256D, and IODP Hole U1309D (Figures 4.2.1.1 and 4.2.1.2; Table 4.2.1.1), have been cored deeper than 1 km into oceanic basement, and these penetrations are arguably some of the greatest technical achievements of scientific ocean drilling. All were “hard won” multiexpedition experiments. From the experiences of drilling these holes, there are important lessons to be learned for the siting, planning, and implementation of future deep drilling of the oceanic basement, which are described in detail in the IODP Expedition 335 Proceedings (Teagle et al., 2012). Establishing the ideal location for drilling is only part of the challenge of successfully drilling moderately deep holes (2–3 km) to recover the samples and data necessary to address long-standing primary goals of scientific ocean drilling. Experience from Holes 504B and 1256D indicates that such experiments require multiple expeditions to achieve their target depths. A total of ∼500 m penetration per expedition is an upper limit for coring in the upper crust, with lesser advances and more frequent drilling challenges as these holes get deeper, and rocks metamorphosed at higher pressures and temperatures are encountered. Penetration and core recovery rates have been low to very low in the two sheeted dike complex sections drilled to date. Average rates of recovery and penetration in the dike section of Hole 1256D are 32% and 0.8 m/h, respectively. The average rate of recovery in the sheeted dike complex of Hole 504B was a miserly 11%. In contrast, experience to date suggests that gabbroic rocks can be cored relatively rapidly at high rates of recovery (e.g., Hole U1309D: penetration rate = 2 m/h; recovery ≥75%), so when the dike–gabbro transition zone is breached in Hole 1256D, solid progress through the plutonic section may be anticipated.
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Scientific ocean drilling has little experience in casing long sections (hundreds of meters) of oceanic basement and a poor armory of underreaming tools for opening hard rock basement holes to the diameters required for the insertion of a casing. Casing hundreds of meters of a deep borehole in igneous basement would be a high risk, costly, and ship-time consuming operation that would produce no new scientific output until completed and drilling was resumed. However, it would greatly improve the stability and hydrodynamics of deep basement boreholes. A regular drilling-then-casing approach to investigate the lower oceanic fast-spread crust (target depth ∼3 km) will require a long-term commitment by the scientific ocean drilling community to a particular site and experiment, and possibly as many as 10 expeditions to complete. The possibility that even such a highly engineered approach could still fail to reach its target would have to be acknowledged and accepted. Such is the capriciousness of hard rock coring that scientific ocean drilling may have to consider new approaches if it is to ever successfully address some of the major science questions that remain unanswered after more than 50 years. There are unlikely to ever be “quick wins” with targets that require multi-expedition deep boreholes. Holes U1309D and Hole 1256D are open, stable, and can be deepened further. Drilling down to the bottom of Hole U1309D was achieved without major difficulty, and the hole condition is excellent. During IODP Expedition 335, Hole 1256D was stabilized, thoroughly cleared of cuttings (several hundred kilograms) and debris. Both holes are ready for deepening in the near future, and fundamental observations concerning the formation and evolution of the ocean crust await.
ACKNOWLEDGMENTS This research used samples and data provided by the Integrated Ocean Drilling Program (IODP). The manuscript benefited from thorough and helpful reviews from Gail Christeson and John Sinton. We thank the USIO teams and JOIDES Resolution crews for their invaluable assistance and outstanding work during IODP Expeditions 304, 305, 309, 312, 335, and 345.
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