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1984 Stratigraphy, Sedimentology, and Geologic Evolution of Eastern Terraba Trough, Southwestern Costa Rica (Tectonics, Forearc Basin). Peter B. Yuan Louisiana State University and Agricultural & Mechanical College
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Recommended Citation Yuan, Peter B., "Stratigraphy, Sedimentology, and Geologic Evolution of Eastern Terraba Trough, Southwestern Costa Rica (Tectonics, Forearc Basin)." (1984). LSU Historical Dissertations and Theses. 4037. https://digitalcommons.lsu.edu/gradschool_disstheses/4037
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Yuan, Peter B.
STRATIGRAPHY, SEDIMENTOLOGY, AND GEOLOGIC EVOLUTION OF EASTERN TERRABA TROUGH, SOUTHWESTERN COSTA RICA
The Louisiana State University and Agricultural and Mechanical Col. Ph.D. 1984
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University Microfilms International
STRATIGRAPHY, SEDIMENTOLOGY, AND GEOLOGIC EVOLUTION OF EASTERN TERRABA TROUGH SOUTHWESTERN COSTA RICA
A Dissertation
Submitted to the Graduate Faculty of the Louisiana State University and Agricultural and Mechanical College in partial fulfillment of the requirements for the degree of Doctor of Philosophy
in
The Department of Geology
by Peter B. Yuan B.S., National Cheng Kung University, 1973 M.S., Carleton University, 1979 December 1984 ACKNOWLEDGEMENTS
It is with deep appreciation that I acknowledge the many people who gave their support to this study. To Dr. Donald R. Lowe, served as chairman of the advisory committee, I give special thanks for his guidance, encouragement, support, and criticisms. Drs. Roy K. Dokka,
Rex H. Pilger, Jr., Barun K. Sen Gupta, Willem A. van den Bold; and
Richard H. Kesel of Department of Geography and Anthropology, served as members of the advisory committee and reviewed the manuscript. Their valuable suggestions and their time and effort are much appreciated.
Many thanks to Dr. Gary R. Byerly who provided assistance in the field and served as member of the advisory committee during the early stage of this study.
To Dr. Stanley H. Frost (Gulf Oil, Exploration, and Production
Company) and Mr. Blake W. McNeely (Shell Oil Company) who identified the foraminifera and provided much needed information on biostratigraphy, I give my sincere thanks.
Many members of the L.S.U. Geology Department have helped in this study. I am grateful for the help and friendship of Patti Phillips,
Antenor Aleman, Gail Fisher, Danny Vforrell, Maud Walsh, Woodson Godfrey,
Arun Majumdar, Ezat Heydari, Barbara Lowery, and Danny Peace. Arthur
Sailer and Amitava Roy performed microprobe analysis of feldspar and pyroxene grains from the study area. Janet Dyson, Michael DiMarco, and
Bruce Nocita read the manuscript and gave many useful suggestions.
During my field work in Costa Rica and Panama, I benefited greatly from Stephen Dixon and William Flores who served as field assistants in
ii the 1979-1981 and 1981-1983 field seasons, respectively. In addition,
Ernesto Araya, Rodolfo Castillo, and the Victor Araya Rodriguez family
of Agua Buena, Coto Brus; Giovanni Lopez of San Vito, Coto Brus; Julian
Gamboa of Florida, Coto Brus; and Manuel Arguello, Enrique Malavassi,
Mr, and Mrs. Tsun Han Wang, and Juan and Carlos Revilla of San Jose,
Costa Rica, are acknowledged for their hospitality and assistance.
To my parents, I owe a special debt of gratitude for taking care
of my children who were away from my wife and me for one year so that we
could spend more time on research. Above all, my greatest debt of
gratitude goes to my wife, Kim, whose patience and understanding proved
to be the best encouragement.
This study was funded by National Science Foundation Grant no. EAR-
78 22754 to Dr. Donald R. Lowe, by grants to Dr. Richard H. Kesel from
the National Geographic Society, and by an Arco Research Fellowship and a Chevron Fellowship. TABLE OF CONTENTS
Page
ACKNOWLEDGEMENTS ...... ii
LIST OF TABLES ...... vi
LIST OF FIGURES...... vii
ABSTRACT ...... ix
INTRODUCTION ...... 1
TERTIARY STRATIGRAPHY OF TERRABA TROUGH ...... 6
SEDIMENTOLOGY OF EASTERN TERRABA TROUGH...... 8
Brito Formation ...... 8
Terraba Formation . . .*...... 10
Lithofacies A: Rhodolite Grainstone and Packstone. .. . 14
Lithofacies B: Cross-stratified Algal-foraminiferal Grainstone...... 16
Lithofacies C: Mudstone and Thin Bedded Sandstone. .. . 24
Lithofacies D: Medium to Very Thick Bedded Sandstone . . 28
Lithofacies E: Cross-bedded Sandstone...... 33
Lithofacies F: Conglomerate...... 35
Relationship Between the Clastic Lithofacies ...... 38
Results of the Markov Chain Analysis and Interpretation. 40
SEDIMENTARY PETROLOGY...... 46
Petrographic Study...... 47
Microprobe Analysis ...... 50
iv TABLE OF CONTENTS (Continued)
Page
TECTONICS AND BASIN EVOLUTION...... 57
Basin Setting ...... 57
Geologic Development...... 59
I - Late Jurassic to Early Eocene...... 59
II - Middle Eocene to Late Eocene ...... 65
III - Oligocene to Early Middle Miocene...... 67
IV - Middle to Late Miocene ...... 74
V - Pliocene to Holocene ...... 75
SUMMARY AND CONCLUSIONS...... 77
REFERENCES ...... 80
APPENDICES
1. Identified Faunas in the Study A r e a ...... 91
2. Markov Chain Analysis of Lithofacies in the Terraba F o r m a t i o n ...... 94
Chi-square T e s t ...... 97
3. Described Stratigraphic Sections...... 98
4. Sandstone Point Count Data ...... 107
VITA ...... 110
v LIST OF TABLES
Table Page
1. Cross-stratification types in Lithofacies B...... 17
2. Subdivisions of Lithofacies D ...... 30
3. Features of large-scale cross-stratifications in Lithofacies E ...... 33
4. Clastic lithofacies of Terraba Formation ...... 39
5. Transitions between sediment gravity flow divisions...... 42
A-l. Lithofacies transition count matrix...... 94
A-2. Lithofacies transition probability matrix ...... 95
A-3. Independent trials probability matrix...... 96
A-4. Difference m atrix...... 96
A-5. Modal analysis of sandstone samples of the Terraba Formation...... 108
A-6. Recalculated values from Table A - 5 ...... 109
vi LIST OF FIGURES
Figure Page
1. Location of Terraba Trough, Nicoya and Osa Peninsulas, and Cordillera de Talamanca...... 2
2. Schematic cross-section showing major tectonic elements of southwestern Costa Rica ...... 2
3. Map of the study area showing locations of samples and measured sections...... 4
4. Geologic map of the study area ...... 5
5. Stratigraphy of Terraba Trough sequence in southwestern Costa Rica and western Panama...... 7
6. Photomicrograph of Brito limestone showing texture and major components ...... 9
7. Four clastic lithofacies of the Terraba Formation...... 12
8. Spheroidal rhodoliths with laminar growth form ...... 15
9. Large-scale cross-stratification in Lithofacies B ...... 18
10. Medium-scale cross-stratification in Lithofacies B ...... 18
11. Small-scale cross-stratification in Lithofacies B ...... 19
12. Paleocurrent directions based on large- and medium- scale cross-stratifications in Lithofacies B ...... 21
13. Flat stratifications in Lithofacies B ...... 22
14. Horizontal trails on bedding plane of rocks in Lithofacies B ...... 22
15. Deformed mudstone in Lithofacies C ...... 26
16. Horizontal burrow Thalassinoides in mudstone of Lithofacies C ...... 27
17. Horizontal burrows Chondrites and Helmlnthoida in mudstone of Lithofacies C ...... 27
18. Representative pattern of bioturbation in mudstone of Lithofacies C ...... 29
19. Lithofacies D 1 ...... 31
vii LIST OF FIGURES (Continued)
Figure Page
20. Lithofacies D 2 ...... 31
21. Lithofacies D 3 ...... 32
22. Lithofacies E ...... 34
23. Mudstone boulders in conglomerate of Lithofacies F ...... 37
24. "Standard" lithofacies sequence of the Terraba Formation based on Markov chain analysis ...... 41
25. Cumulative thickness distribution for all sandstone layers of the Terraba Formation...... 45
26. Q-F-L diagram...... 48
27. Anorthite content of detrital plagioclase grains ...... 51
28. Map showing maximum anorthite content (mol %) of detrital plagioclase grains at representative locations ...... 52
29. Pyroxene quadrilateral ...... 53
30. Plot of titanium versus aluminum for detrital pyroxene g r a i n s ...... 55
31. Correlation between major tectonic events and evolution of Terraba Trough ...... 60
32. Distribution of "Nicoya Complex" and Eocene shallow-marine limestone in Costa Rica, Panama and Colombia ...... 62
A-l. Map showing location of fossil samples listed in Appendix 1 ...... 90 ABSTRACT
Sedimentary rocks of eastern Terraba Trough, southwestern Costa
Rica, were deposited in a fore-arc basin developed at an ocean-ocean convergent boundary. The basin developed in Middle to Late Eocene when the Farallon Plate began its subduction beneath the Caribbean Plate.
Shallow-water carbonates of the Brito Formation were deposited on shoals of uplifted basement blocks. These were surrounded by deeper-marine areas in which mixed volcaniclastics and carbonate debris accumulated.
The Brito Formation consists of algal-foraminiferal packstone to grain stone, rudstone, and rare wackestone formed in fore-slope, carbonate buildup, and open platform environments in a warm, tropical sea.
The environmental conditions remained essentially unchanged during the Oligocene. Within the study area, the Eocene Brito Formation is overlain directly by rocks of the Upper Oligocene Rio Claro Member of the Terraba Formation, a newly defined carbonate unit. It is composed of rhodolite and bioclastic grainstone deposited in a minimum water depth of 6 meters. A combination of relatively little subsidence, mild volcanism to produce only small amount of clastic debris, and possible erosion at about 30 Ma during a global drop of sea level may be responsible for the absence of Lower Oligocene rocks in the study area.
After the deposition of the Rio Claro Member, the study area sub sided rapidly to become a trough possibly deeper than 2,000 meters.
Trace fossils, the association of sedimentary structures, and Markov chain analysis of clastic lithofacies indicate that sedimentation occurred in deep water from sediment gravity flows.
ix In Early to early Middle Miocene, coarser sediments and thicker
sand units containing coal fragments, Pecten, and Turritella became more
abundant, suggesting that the basin was gradually filled.
Petrographic study and microprobe analysis of sediments indicate
that the source for the bulk of the detrital material in the Terraba
Trough was a rather homogeneous, basaltic andesite to basaltic magmatic
arc characterized by an explosive eruptive style. Earlier appearance of
rocks with alkaline affinity in the study area than in the central and western Terraba Trough suggests that the study area was close to a hinge
fault or in an area with slow subduction.
The results of this study indicate that the timing and degree of
subsidence of the fore-arc basin, the vertical variations in lithology, and the geologic evolution of eastern Terraba Trough are directly related to the Mesozoic and Tertiary tectonic evolution, in particular, the variations in convergence rate between lithospheric plates in this part of Central America and the eastern Pacific Ocean.
x INTRODUCTION
Geologically, Central America can be subdivided into two distinct provinces. Northern Nicaragua and areas to the north are characterized by a continental crust and pre-Pennsylvanian metamorphic basement under lying upper Paleozoic and younger rocks. In contrast, southern Nicara gua, Costa Rica, and Panama have a Jurassic to Cretaceous oceanic base ment underlying Cenozoic strata (Weyl, 1980).
In Costa Rica, Mesozoic basement rocks are exposed only along the
Pacific coast, most prominently on the Nicoya and Osa Peninsulas
(Fig. 1). This sequence, called the Nicoya Complex (Dengo, 1962), con sists of highly deformed pillow basalt, diabase, ultramafic rock, gray- wacke, mudstone, carbonate* and chert. It is inferred to represent an uplifted segment of late Mesozoic oceanic crust that has been accreted along the Middle America Trench (Stibane et al., 1977; Galli-Oliver,
1979; Schmidt-Effing, 1979). The Nicoya Complex on the Osa Peninsula is in fault contact with and possibly overlain by lower Tertiary sedimen tary rocks of the Terraba Trough. The Trough is a fault-bounded Quater nary structural depression containing a synclinally deformed Eocene to
Holocene fore-arc basin sequence. It is located between the Cordillera de Talamanca, a segment of the Middle America magmatic arc, and the off shore Middle America Trench (Fig. 2). Except for brief reconnaissance by Henningsen (1966a, b), little geological work has been published on this area.
The utilization of the petrology and sedimentology of fore-arc basin sequences to document the tectonic and magmatic evolution of con- 83' NICARAGUA
NICOYA COMPLEX
50 km
o NICOYA
PANAMA MEXICO
OSA NICARAGUA COSTA RlCA
PANAMA 83°
Fig. 1. Location of Terraba Trough (Tertiary fore-arc basin se quence), Nicoya and Osa Peninsulas (Mesozoic basement complex), and Cordillera de Talamanca (magmatic arc assemblage). Cross-section along X-X' is shown in Figure 2.
CORDILLERA DE TALAMANCA TERRABA OSA PENINSULA TROUGH
PACIFIC CARIBBEAN
NICOYA SUBDUCTION / FOREARC BASIN MAGMATIC ARC BACKARC BASIN OBDUCTION COMPLEX SEQUENCE ASSEMBLAGE SEQUENCE
M . A M . TRENCH
Fig. 2. Schematic cross-section along X-X' of Figure 1, showing major tectonic elements. Distance of X-X’ is twice as that in Figure 1. (Modified from Dengo, 1962). vergent boundaries has received considerable attention in recent years
(Dickinson and Seely, 1979; Dickinson, 1982; Leggett, 1982; Hayes,
1983). To date, these studies have dealt mostly with fore-arc basins developed in association with Andean-type margins. Studies of ocean- ocean fore-arc basin sequences are lacking, largely because such se quences are poorly represented in the geologic record. They are gener ally strongly metamorphosed and tectonized by later orogenic events.
The largely unmetamorphosed Terraba Trough sequence of southwestern
Costa Rica offers an excellent opportunity for studying the stratigra phy, sedimentology, and geologic evolution of a fore-arc basin sequence developed at an ocean-ocean convergent boundary.
The objectives of this study are: (1) to document the onshore geology of the Terraba Trough in southeastern-most Costa Rica and west ern Panama (Figs. 3, 4) as an example of a fore-arc basin developed at an ocean-ocean convergent boundary, and (2) to utilize this geology together with the results of offshore research (Handschumacher, 1976;
Hey, 1977; Hey et al., 1977; Lonsdale and Klitgord, 1978), to interpret the Tertiary history and geologic evolution of this portion of Central
America. COSTA .fp RICA
STUDY AREA
,N VITO
1300
75 • AGUA 8*45- 57 BUENA
100 m '0 m <■67 66
-8*40'
INTERAMERICAN
VILLA V-33 NEILY) O
Km 82*55'
Fig. 3. Map of the study area in southwestern Costa Rica showing topography and locations of samples and measured sections. 5
J 22 ill 32^ALTO H O O O m ^ GARROTE
H600m /v /W\
f-200m BRITO®.\\\\JERRABA FM. »\>\\\\\\\\\'
KM
sCTTO 111111111I I i I MK\I W vA \ v
LEGEND
Fig. 4. Geologic map and cross-section of the study area. The Curre Formation (not shown) is probably covered by the Pliocene Paso Real Formation (?) and Quaternary Brujo Formation. Also shown are strikes and dips of sandstone beds of the Terraba Formation.
Distribution of Paso Real and Brujo Formations partly based on Henningsen (1966a) and D. R. Lowe (personal commun.). TERTIARY STRATIGRAPHY OF TERRABA TROUGH
The Tertiary sequence in the Terraba Trough of southwestern Costa
Rica includes more than 4500 m of strata ranging in age from Middle
Eocene to Holocene. Dengo (1960, 1962) and Henningsen (1966a) subdi vided the strata into four formations (Fig. 5). D. R. Lowe and R. H.
Kesel (personal commun.) have further recognized an as yet unnamed
Pliocene mudstone unit and the Quaternary Brujo Formation.
The overall succession in the Terraba Trough represents a major transgressive-regressive cycle (Phillips, 1983). Middle to Late Eocene shallow-marine carbonate sedimentation (Brito Formation) was followed by transgression. From Oligocene to Early Miocene time, sedimentation occurred predominately under deep-water conditions (mudstone and sand stone of the Terraba Formation). The upper Terraba Formation and the succeeding Curre Formation reflect shoaling from fully subaqueous through shallow-water to largely subaerial conditions (Lowery, 1982),
The as yet unnamed Pliocene mudstone reflects marine transgression
(D. R. Lowe and R. H. Kesel, personal commun.), which was followed by the deposition of terrestrial volcanic and volcaniclastic rocks of the
Pliocene Paso Real Formation and Quaternary non-marine sands and gravels of the Brujo Formation. Quaternary deposition was associated with uplift of the Cordillera de Talamanca and deformation of the fore-arc basin to form the Terraba Trough (D. R. Lowe and R. H. Kesel, personal commun.).
6 7
Central and Western Terraba Trough STUDY AREA (Eastern Terraba Trough)
AGE (M a) HOLOCENE Brujo Fm. 0. 01- PLEISTOCENE PLIOCENE 5.3 - Gisrr© Fm, 11.2- MIOCENE 16. 6- T e rra b a T e rra b a 2 3. 7- Fm. Rio Claro Fm. M em ber 3 0 . 0 - OLIGOCENE
3 6. 6-
4 0 . 0 - Fm. Brito Fm.Brito L EOCENE 5 2. 0 -
57.8
PALEOCENE
66 .4
< » N icoya uiO CCLU oo Com plex 144
208
Fig. 5. Stratigraphy of Terraba Trough sequence in southwestern Costa Rica and western Panama. Stratigraphy of central and western Terraba Trough modified from Hutchison (1980) and Weyl (1980).
* Symbol represents the position of the unnamed Pliocene mudstone unit (D. R. Lowe and R. H. Kesel, personal commun.).
Absolute time scale after Palmer (1983). SEDIMENTOLOGY OF EASTER LERRABA TROUGH
Brito Formation
The oldest rocks recognized in the Terraba Trough are carbonates of
the Middle to Upper Eocene Brito Formation. The basal contact of the
Brito Formation with the Nicoya Complex is not exposed but is apparently either an unconformity or a fault (Dengo, 1962). The formation has a thickness of 1000 m (Henningsen, 1966a) and, in the study area, forms continuous high-standing ridges made up of massive, white to yellowish gray, fossiliferous limestone. According to Henningsen (1966a), in ad dition to limestone, the Brito Formation includes tuffaceous sandstone, conglomerate, marl, and mudstone. These lithotypes, however, were not found in the Brito within the study area.
In all outcrops where it was observed, the Brito is uniform in tex ture and composition. It is composed of algal-foraminiferal packstone to grainstone, rudstone, and rare wackestone (classification after
Dunham, 1962; Embry and Klovan, 1971). Texturally, the Brito samples are most similar to the Standard Microfacies (SMF) types 5, 6, 10, and
18 of Fltigel (1972, 1982) and Wilson (1975).
Algal and foraminiferal debris ranges in size from fine silt to sand to particles up to 3 cm in diameter and is invariably poorly sorted (Fig. 6). The algae are mostly broken, commonly abraded, and include both branching and encrusting coralline algae and some dasycladacean green algae. The foraminifera are mostly larger benthic forms, up to 2 cm in length (Appendix 1). Other allochems include, in
8 Fig. 6. Photomicrograph of Brito limestone, showing low degree of sorting, presence of lime mud, and major components: algal fragments (dark grains), larger foraminifera, pellets, and echinoid (bottom center). Length of foraminifera at center =2.5 mm. 10
order of decreasing abundance, pellets and peloids; debris of echinoids, bivalves, and bryozoans; and planktic foraminifera.
The assemblage of larger benthic foraminifera, e.g., Lepidocyclina and Discocyclina, coralline and green algae, echinoids, bivalves, and bryozoans indicates that the Brito Formation was deposited in a shallow, warm, tropical sea with open circulation and normal salinity (Ross,
1979; Wilson and Jordan, 1983). The light color of the limestone also suggests that the allochems were deposited in we11-oxygenated water, probably above storm wave base (Wilson, 1975).
The identified SMF (Standard Microfacies) types can be best assign ed to the Standard Facies Belts (Wilson, 1975) 4, 5, and 7 representing foreslope, carbonate buildup, and open platform environments, respec tively. The wackestone was deposited in localized low areas on the mid shelf such as the mud-filled Holocene Hawk Channel behind the Florida
Reef Track (Wilson and Jordan, 1983), whereas the packstone/grainstone formed on shelf shoals or on flanks of carbonate buildups.
Terraba Formation
The Brito Formation is overlain by a thick sequence of Lower
Oligocene to lower Middle Miocene clastic rocks named the Terraba Forma tion by Henningsen (1966a). Within the study area, it includes 1200-
2000 m of flysch-like interbedded mudstone, volcaniclastic sandstone, and minor conglomerate (Dengo, 1960; Henningsen, 1966a; Phillips, 1983).
In the study area, rocks of this lithology occur both stratigraph- ically above (north) and structurally below (south) of outcrops of the
Brito Formation (Fig. 4). Foraminifera collected for this study indicate that the formation ranges in age from Late Oligocene to early
Middle Miocene (S. H. Frost, personal commun.; B. W. McNeely, personal commun.). There is strong evidence that the Lower Oligocene is missing in the study area. Close sampling at three locations (PY-63, 74, and
75), where the contact between the Brito and Terraba Formations is well exposed, shows that Upper Oligocene rocks directly overlie Middle to
Upper Eocene carbonates. There is no indication of faulting at the contact, so it is unlikely that the Lower Oligocene has been faulted out. It seems most probable that the contact is an unconformity.
For the purpose of sedimentological analysis, the Terraba Formation in the study area is here divided into six lithofacies, each character ized by distinctive lithology and association of sedimentary structures.
Lithofacies A: Rhodolite grainstone and packstone Lithofacies B: Cross-stratified algal-foraminiferal grainstone Lithofacies C: Mudstone and thin bedded sandstone Lithofacies D: Medium to very thick bedded sandstone, containing three subfacies: Dl, D2, and D3 Lithofacies E: Cross-bedded sandstone Lithofacies F: Conglomerate
Lithofacies A and B are carbonate rocks of Oligocene age, with B overlying A. Except for one location (Henningsen, 1966a), no Oligocene limestone has been reported previously in the Terraba Trough. The name
Rio Claro Member is here assigned to these newly discovered lithofacies.
The type locality is in the Rio Claro (8045,21", 83 0 0' 5 8 !1)_, jwhere Litho facies A and B are well exposed. The four clastic lithofacies C, D, E,
F are illustrated graphically in Figure 7. Fig. 7. The four clastic lithofacies of the Terraba Formation illustrated by using selected intervals from measured sections.
Lithofacies C is from section number PY-33. Lithofacies D : PY-78 and PY-77. Lithofacies E : PY-71. Lithofacies F : PY-35. IT
very c sandstone mudstone boulders matrix-supported
vary c sandstone D3 M assive 10' medium to e pebbly conglomerate sa n d sto n e S3 sa n d y S 3 02 coarse to v coarse S2 Stratified to graded s a n d sto n e -25 vc sandstone -21 coarse to v coarse S2 very c sandstone s a n d s to n e 01 a s h la y e r ■
Amalgamated med to c sandstone 24 mudstone -20 v fine sandy gastropods near top med to coarse s a n d s to n e mudstone b o u ld ers up to 70 cm long -23 -7 n '19m fine to med sandstone lens pebbly conglomerate san d y fine sandstone
22 r
Lithofacies C Lithofacies D Lithofacies E Lithofacies F Mudstone and M ed iu m t o Very Thin Bedded Sandstone T h ic k Bedded S a n d s t o n e Cross-bedded Sandstone Conglomerate 14
Lithofacies A: Rhodolite Grainstone and Packstone
This lithofacies overlies Middle to Upper Eocene rocks of the Brito
Formation with sharp contact. Although only 3.5 m thick, it can be traced for at least 21 km along strike (Fig. 4). The rock consists of rhodoliths in a matrix of bioclasts.
The rhodoliths range in diameter from 0.5 to 4 cm, averaging
2.5 cm, and are moderately to well sorted. Most are spheroidal, some ellipsoidal, and a few discoidal (terminology of Bosence, 1983a).
Cross-sections of the rhodoliths show coralline algae nucleating about algal fragments, coral, or pebbles of volcanic rock (Fig. 8). The growth forms of the coralline algae around the nuclei are predominantly laminar, i.e., concentrically and continuously laminated, although some are locally columnar (Bosence, 1983a). Larger rhodoliths are more heavily bioeroded and have more crenulate surfaces than smaller ones.
The sand-sized matrix consists of debris of coralline algae and echi noids, and abundant larger benthic and encrusting foraminifera.
Modern rhodoliths are found from the lowest intertidal zone down to
90 m but are most abundant between depths of 1 and 10 m (Bosence,
1983b). The coral and echinoid fragments and larger benthic foraminif- era associated with the rhodoliths also indicate that these rhodoliths have a shallow-marine origin.
Although a complex set of physical, chemical, and biological factors controls the formation of rhodoliths, the primary factor is current or wave action. Strong water motion prevents the fragile rhodoliths from forming, whereas weak wave or current action leads to algae stablization through growth and coalescence or to burial by sedi- Fig. 8. Spheroidal rhodoliths with laminar growth form. Those with thinner walls (center left and right) are more irregular in shape. Pen at top for scale. 16
ments that eventually kill the algae (Adey and Macintyre, 1973) „
Periodic but infrequent turn-around by water is necessary in order to produce the nodular form.
Based on study of living rhodoliths, Besellini and Ginsburg (1971) noticed the close relationship between the algae growth forms (branch ing, globular, columnar, and laminar) and the frequency of rhodolith movements by store-generated currents. Their study indicates that rhodoliths with spheroidal shape and laminar growth form, such as those characterizing Lithofacies A, occur in environments with the highest frequency of movement, such as channels. Continued growth and frequent movement produce successive smooth coatings. As the rhodoliths grow larger, they become less mobile and are likely to be trapped in depres sions and subject to bioerosion and burial. Growth of algae may contin ue on the immobile rhodoliths for some time, but the surface becomes crenulate due to lack of smoothing through rolling.
A shallow, tropical carbonate shelf, probably less than 10 m deep and periodically swept by storm currents, is the most probable environ ment in which sediments of Lithofacies A were deposited.
Lithofacies B: Cross-stratified Algal-foraminiferal Grainstone
Overlying the rhodolites of Lithofacies A is a 10 m-thick unit of carbonate grainstone here termed Lithofacies B. Based on the identified larger foraminifera, this lithofacies is Late Oligocene in age (S. H.
Frost, personal commun.) and has a lateral extent of at least 4 km. It grades into the underlying rhodolite through a 1 m-thick transition zone, and upward into the overlying mudstone or silty to very-fine-sandy 17
claystone.
The rock is composed of well-sorted bioclasts, mostly coarse- to
very-coarse-sand-sized, but ranging from medium sand to granules. The
bioclasts are coralline algae (45%), larger benthic foraminifera (40%),
echinoids (15%), bryozoans (2%), planktic foraminifera (2%), and
mollusks (1%), with traces of glauconite and plagioclase. Lime mud is
rare to absent. Rock color varies from light olive gray to greenish
gray: an abundance of foraminifera makes the rock light olive gray; whereas it appears greenish gray where the algal content is high.
Lithofacies B is characterized by the predominance of cross
stratification. Based on their thickness, the cross-sets can be sub divided into three types (Table 1 and Figs. 9, 10, 11).
Table 1. Cross-stratification types in Lithofacies B.
Type Cross-set Cross-set Thickness Foreset Foreset dip thickness geometry of foreset geometry (degrees) layers
Large-scale 0.6-1 m Tabular 5-10 cm Tangential to 12-21 sightly concave upward
Medium-scale 0.2—0.5 m Tabular 2.5-5 cm Angular to 5-12 tangential
Small-scale <0.1 m Tabular 1-2 cm Angular to 10-13 tangential
Number of observations of large-scale cross-sets: 3; medium-scale cross-sets: 12; small-scale cross-sets: 6.
The following features are common to all cross-sets:
(1) They form tabular sets: within the limit of the exposures (less than
5 m), there is no observable convergence or curvature of the bound- 18
Fig. 9. Large-scale, tabular cross-stratification with tangential to slightly concave-upward foresets in Lithofacies B.
Fig. 10. Coset of medium-scale tabular cross-stratification in Litho facies B. 19
Fig. 11. Coset composed of four small-scale tabular cross-sets in Lithofacies B. ing surfaces.
(2) The dip angle of the foresets is low: they range from 5 to 21
degrees, but most are near 10 degrees.
(3) There is no apparent change in grain size along the dip of the
foresets.
(4) Paleocurrent directions, based on the foreset orientation, are
confined within a span of 124 degrees, and are not highly diversi
fied (Fig. 12).
Other sedimentary structures include flat stratification composed
of alternating dark- and light-colored layers. The dark layers are 1-2
cm thick and composed predominantly of medium- to coarse-sand-sized
algal fragments; whereas the light layers are 2-4 cm thick and made of
abundant foraminifera ranging in length from 0.5 to 4 mm (Fig. 13).
Structureless beds of limestone also occur. On bedding planes, large horizontal trails are not uncommon (Fig. 14).
The high degree of sorting of the allochems and absence of lime mud
throughout this lithofacies indicate sediment deposition in an environ ment dominated by constant current and/or wave activity, probably on a carbonate shelf. The shelf environment is also indicated by the predom inance of shallow-marine fossils. The thickness of the cross-sets and their tabular shape suggest that the bedforms were sand waves with straight to slightly sinuous crests (Harms et al., 1982). Had the flow strength been higher, dunes with festoon cross-sets would have been generated.
Similar tabular cross-stratifications in modern carbonate shelf sands with foreset slopes less than the angle of repose have been termed 21
N4°E
N 64 W ,N60 E
Fig. 12. Paleocurrent directions based on orientations of foresets of three large- (with L symbol) and five medium-scale cross-sets in Lithofacies B. Fig. 13. Flat stratifications (middle) over- and underlain by small-scale tabular cross-sets. Apparent current direction is from right to left.
Fig. 14. Horizontal trails on bedding plane of algal-foraminiferal limestone of Lithofacies B. 23
"accretion ripples" (Imbrie and Buchanan, 1965). Conditions favoring accretion ripple formation are considered to be high flow velocity and water depth that is large compared to the bottom relief. Flume experi ments by Jopling (1965) with fine to coarse sands support these inter pretations. A relationship appears to exist between current velocity and foreset slope angle. On the Bahama Platform, where the mean current velocity is high (up to 2 m/sec), both high- and low-angle tabular cross-sets are present (Imbrie and Buchanan, 1965; Hailey et al., 1983), whereas on the Yucatan Shelf, where current velocity is generally less than 0.5 m/sec, the sand wave foreset slopes are very close to the angle of repose (Harms et al., 1974).
The hydraulic conditions under which Lithofacies B was formed can be interpreted as follows:
(1) The water depth was considerably greater than the height of the sand
waves. Also, the range of water level fluctuation was small
relative to water depth. Relatively deep water not only prevented
the formation of dunes, which requires a higher flow strength, but
also provided favorable conditions for forming low-angle foresets.
Field, analytical, and flume studies have shown that at depths
greater than several tens of cm the mean height of relatively
straight-crested sand waves is generally less than or equal to
approximately 1/6 of the mean flow depth (Rubin and McCulloh, 1980).
The minimum water depth for generating sand waves of this litho
facies is therefore estimated to be 6 m.
The maximum water depth is not as evident. Studies on the
Bahama Platform (Imbrie and Buchanan, 1965; Ball, 1967; Hailey et
al., 1983) and Yucatan Shelf (Harms et al., 1974) indicate that the predominant bedforms are sand waves with straight to slightly
sinuous crests, all occurring in water depth less than 6 m.
Elsewhere, similar bed forms in non-carbonate sediments bave been
reported in water depths greater than 100 m (Bouma et al., 1980;
Hamilton et al., 1980). However, the large grain size (medium-
grained sand to granules) and dominance of shallow water allochems
within this lithofacies suggest a relatively shallow depth.
(2) The differences in the three recognized cross-stratification types
(Table 1) can probably be related to changes in current velocity.
Rubin and McCulloch (1980) have shown that at a constant depth, sand
wave height increases as flow velocity increases. When the current
velocity was lower than required for generating larger bedforms, the
carbonate sands formed either lower plane beds or were subject to
bioturbation.
Lithofacies C: Mudstone and Thin Bedded Sandstone
This lithofacies, making up most of the Terraba Formation in the study area, consists of mudstone and thin bedded sandstone. The sand stone/mudstone ratio ranges from 1:5 to 1:14.
The sandstone is grayish olive green to dark greenish gray and generally medium- to very-coarse-grained, but ranges from very-fine grained sandstone to granule conglomerate. Unless disrupted by bio
turbation, the sandstone beds are laterally continuous and range in thickness from 1 to 60 cm.
Sandstone beds display three sequences of sedimentary structures:
(1) Normally graded massive sandstone overlain by flat laminated sandstone.
(2) Flat laminated sandstone throughout.
(3) Flat laminated sandstone overlain by festoon cross-laminated
sandstone.
The mudstone is dark greenish gray and up to 15 m in thickness.
It locally shows evidence of soft-sediment deformation, with the de formed zones generally 50-60 cm thick and traceable for up to tens of m
(Fig. 15). In most places, the mudstone contains admixed very fine to fine sand. Thin stringers of fine to medium sands are commonly inter laminated within the mudstone. Bioturbation is extensive, suggesting that much of the sandy mudstone was originally clay-rich mud containing thin sand interbeds. These lithologies were later homogenized by burrowing organisms.
Burrows are essentially horizontal and occur in three different sizes and shapes :
(1) Thalassinoides: branched burrows without surface ornamentation,
1.5-2 cm in diameter and up to 15 cm in length, forming complex
criss-cross patterns (Fig. 16).
(2) Chondrites: branched burrows in which branches are 1 cm in diameter
and up to 10 cm in length. Each branch diverges at approximately 45
degrees from the previous branch (Fig. 17).
(3) Helminthoida; fecal ribbons 1 mm in diameter and up to 1 cm in
length (Fig. 17). This is the most abundant trace fossil within the
mudstone.
The trace fossils Thalassinoides, Chondrites, and Helminthoida, are typically associated with deep-water (outer shelf to bathyal) sedi- 26
Fig. 15. Deformed mudstone in Lithofacies C. 27
Fig. 16. Horizontal burrow Thalassinoides in mudstone of Litho facies C.
Fig. 17. Two more types of horizontal burrows: Chondrites at right showing 45 degrees diverging and equal-width branches; Helminthoida at left (light and dark fecal ribbons). 28
ments (Chamberlain, 1978). The predominance of mudstone in this litho
facies and the high intensity of horizontal bioturbation indicate that
the fine, suspended sediment accumulated in quiet water at sufficiently
low sedimentation rates that organisms had adequate time to rework the
deposits. However, the slow but continuous sedimentation of fine sedi
ments was punctuated by rapid, high-energy depositional events repre
sented by interbedded coarse-grained, largely non-bioturbated, and
current-deposited sandstone layers (Fig. 18).
The high lateral continuity of the thin bedded sandstone, together
with its characteristic sedimentary structures, suggests that the sands
were deposited by turbidity currents. The graded bedding, flat lami
nation, and cross-lamination represent Bouma Ta, Tb, and Tc divisions,
respectively, and were deposited by "low-density" turbidity currents
(Bouma, 1962; Lowe, 1982). Those beds containing coarse sand or coarser
sediments, however, may have been deposited by "high-density" turbidity
currents (Lowe, 1982). Local instability of the slope on which sedimen
tation took place resulted in downslope movement of semi-consolidated
sediment and formation of the soft-sediment deformation structures.
Lithofacies D: Medium to Very Thick Bedded Sandstone
This lithofacies is composed of medium (10-30 cm) to very thick
bedded (greater than 100 cm) sandstone interbedded with mudstone. It
makes up 35% of the measured sections. The sand/shale ratio is greater
than 1:2.
Based on differences in bed thickness and sedimentary structures,
Lithofacies D can be divided into three subfacies : Dl, D2, and D3 Fig. 18. Lithofacies C: mudstone (light) and sandstone (dark).
The mudstone was totally burrowed as indicated by its mottled appear ance and the presence of abundant trace fossils.
Thinner sandstone bed (dark layer about 1/4 from bottom) was more heavily biotubated than the thicker sandstone bed (dark layer under the pen) so that the thinner sandstone is no longer traceable laterally. The thicker sandstone bed was burrowed more intensely at its top than bottom. (Table 2).
Table 2. Subdivisions of Lithofacies D.
Dominant Bed Sedimentary Sediment flow Lithofacies grain size thickness structure division *
Dl medium to up to 100 cm, stratification, S2, Tt (Fig. 19) coarse formed by some reverse amalgamation grading of layers <10 cm thick
D2 medium to up to 600 cm, stratification S2, S3 (Fig. 20) very coarse formed by to amalgamation normal grading of layers >10 cm thick
D3 medium to up to 400 cm, massive, some S3 (Fig. 21) very coarse most <100 cm normal grading
* Based on Lowe (1982).
Bioturbation is less intense than in Lithofacies C. Horizontal
Thalassinoides are the predominant burrows found. Carbonaceous
flakes several mm long and coal fragments up to 5 cm by 12 cm in size
are locally abundant.
Characteristic sedimentary structures (Table 2) and their vertical
succession in sandstones of this lithofacies suggest deposition by sandy high-density turbidity currents. The amalgamation of thin sandstone
layers into thick beds (Table 2) is common in sediment gravity flows
(Walker and Mutti, 1973). The flows derived part of their sediments from a land area which sustained vascular plants as suggested by the coal fragments present in this lithofacies. Fig. 19. Lithofacies Dl: Interbedded medium to coarse-grained sand stone (light) and thin bedded mudstone (dark). The sandstone beds are formed by amalgamation of layers < 10 cm thick.
Fig. 20. Lithofacies D2: Stratified to normally graded, medium- to very coarse-grained sandstone. The sandstone bed thickness increases u pw a r d s . Fig. 21. Lithofacies D3: Massive, medium to very coarse-grained sandstone overlying mudstone along wavy boundary, possibly reflecting loading. Pen cap at lower left for scale. 33
Lithofacies E: Cross-bedded Sandstone
Cross-sets up to a few cm in thickness are common in the thin
bedded sandstone of Lithofacies C. Rocks containing cross-stratifica
tion of a much larger scale, with sets up to 1 m thick, are grouped
separately into Lithofacies E. This lithofacies makes up about 1% of
that of the sections measured. It consists of well-sorted, granular
very coarse-grained sandstone. Features of the cross-stratification are
listed in Table 3.
Table 3. Features of larg e-scale cross-stratification in Lithofacies E.
Cross-set Cross-set Thickness of Foreset Apparent foreset thickness geometry foreset layers geometry dip (degrees)
30 cm to trough to 2. 5-12 cm concave-up 25-30 1 m tabular ? tangential
Number of observations = 5.
All cross-stratified units are in sharp contact with the underlying mudstone or stratified, fine- to coarse-grained sandstone. The cross sets grade upward into massive or horizontally stratified fine- to coarse-grained sandstone (Fig. 22).
Large-scale trough cross-stratification with high-angle, concave- upward foresets is generally formed by the migration of 3-dimensional dunes (Harms et al., 1982). The lower sharp contact of the cross-sets suggests scouring of underlying sediments before deposition of foreset sediments, a feature characteristic of migrating dunes (Harms et al.,
1982). The massive or stratified sandstones overlying the large-scale cross-stratification were probably produced by fall-out or traction Fig. 22. Lithofacies E: Very coarse-grained sandstone containing large-scale cross-stratification overlying mudstone along sharp contact (hammer). The high-angle foresets, running from upper right to lower left, have tangential bases. sedimentation, respectively. Vertical succession of the above struc
tures therefore indicates change of environmental conditions from (1)
slow sedimentation (mudstone) through (2) scouring and deposition by
traction (cross-stratification with sharp base) to (3) traction carpet
or suspension sedimentation, an assemblage close to that produced by
high-density turbidity currents (Lowe, 1982).
Hiscott and Middleton (1979) reported similar large-scale cross
stratification and vertical structural transitions in Ordovician sand
stones deposited at the base of a submarine slope. In the study area, other clastic lithofacies (C, D, F) which are vertically and laterally
associated with Lithofacies E have been interpreted to be deposits of sediment gravity flows. Thus, based on the vertical succession of
sedimentary structures and lithofacies association, Lithofacies E most likely represents deposits of sediment gravity flows in relatively deep water.
Lithofacies F: Conglomerate
Conglomerate (some breccia) accounts for 3 percent of the total
thickness of measured sections. It is predominant only in one quarry
section (PY-35) from which the following description is taken. At the
quarry, the conglomerate consists of clast-supported, moderately sorted,
subangular to rounded, volcaniclastic debris ranging from coarse sand to gravel 2 cm in diameter. Coal fragments up to 7 cm in length, broken gastropod and bivalve shells, and mudstone chips are locally abundant.
Five conglomerate beds, 0.4 m, 5.7 m, 0.2 m, 2.8 m, 3.1 m in thick ness, can be identified at the quarry. They are crudely stratified 36
internally and in places contain reversely graded zones at their bases.
Near their tops, most of these conglomerate beds contain boulder zones.
The zones are 20 cm to 2 m thick and composed of well-rounded, matrix- supported boulders up to 70 cm long. The boulders show little or no grading or imbrication (Fig. 23). All boulders are composed of mudstone and some contain foraminifera and bivalves.
Interbedded with the conglomerate are mudstone and sandstone layers.
The mudstone layers are horizontally burrowed. In the quarry section, a 20 cm by 5 m mudstone block was found to have been recumbently folded before lithification and then embedded in sands.
The planktic foraminifera Globigerinoides sicanus, Globorotalia mayeri, and Globorotalia peripheroronda (B. W. McNeely, personal commun.) found in the sequence demonstrate deposition in marine environ ment. Deposition of the conglomerate and the interbedded mudstone, however, required contrasting mechanisms of transportation and sediment ation.
The mudstone represents slow sedimentation of fine sediments through settling in a low-energy, probably deeper marine environment.
The conglomerate represents "resedimented" volcanic rock fragments which were originally deposited in a shallow-marine, near-shore environment as suggested by the associated gastropod, bivalve, and coal fragments.
The clast-supported texture and stratification of the sandstone and conglomerate suggest that it was deposited by sediment gravity flows possibly ranging from gravelly high-density turbidity currents to debris flows (Lowe, 1976, 1982). On their way downslope, these flows eroded the underlying mud, incorporating large chunks into the flow. Similar processes have been observed in modern submarine canyons (Shepard and 37
Fig. 23. Mudstone boulders near top of conglomerate bed (Lithofacies F). The boulders are matrix-supported and show little or no imbrication. Dill, 1966; Palmer, 1976).
Relationship Between the Clastic Lithofacies
Trace fossils and benthic foraminifera (Appendix 1) indicate that
Lithofacies C, D, E, and F were deposited under deep-water bathyal conditions. The presence of Melonis pompilioides in sample #74-4 (Upper
Oligocene) and #73-1, #73-2, #73-3 (Lower Miocene to lower Middle
Miocene), covering the total time span of the Terraba Formation in the study area, indicates a possible water depth greater than 2000 m (Bandy and Chierici, 1966) during the deposition of the Terraba clastic litho facies. The use of Melonis pompilioides as a lower bathyal depth in dicator, however, has been questioned; the species apparently has changed its depth-adaptation in the course of its geological history
(Blake and Douglas, 1980). The sedimentary structure suite in sand stones of these lithofacies indicate deposition by sediment gravity flows. The lithologies, sedimentary structures, and sediment gravity flow divisions of the Terraba clastic lithofacies are summarized in
Table 4. 39
Table 4. Clastic lithofacies of Terraba Formation.
Total % of Litho Litho- Sedimentary Sediment flow thickness of all facies i°gy struc tures division * measured sections
C mudstone and normal grading, Tt-e, S3 61% thin bedded stratification, sandstone cross-lamination
D1 medium bedded stratification, S2, Tt 13% sandstone some reverse grading
D2 very thick stratification S2, S3 14% bedded to normal grading sandstone
D3 very thick massive, some S3 8% bedded normal grading sandstone
E thick bedded large-scale SI 1% sandstone cross-bedding
F conglomerate stratification, R 2 , R3 3% reverse grading
* Based on Lowe (1982).
It is critical to establish the vertical and lateral relationship between the Terraba clastic lithofacies C, D, E, and F for overall envi ronmental interpretation. In the study area, however, rarely did all four lithofacies occur together or show consistent vertical relation ships. The correlation of lithofacies between outcrops was also extremely uncertain because of poor exposure and heavy vegetation.
Markov chain analysis was therefore applied to all measured sections in order to identify any preferred stratigraphic arrangement of the litho facies. Interpretation can then be made on the evolution of deposition- al mechanisms in the study area. Furthermore, a chi-square test was 40
applied to evaluate the significance of the results of the Markov chain
analysis. Chi-square test and Markov chain analysis procedures are
included in Appendix 2.
Results of the Markov Chain Analysis and Interpretation
The results of the Markov chain analysis are presented in Table 5 and Figure 24. Several observations can be drawn from these results:
(1) The most common sequence begins with Lithofacies C at bottom grading
upwards into D3. Above D3, there is likely to be an interbedding of
D3 and F.
(2) D1 and E have a tendency to occur interbedded with one another, as do
D2 and F.
(3) F is more likely to underlie D2 and D3 than to overlie them.
Similarly, E is more likely to underlie than to overlie Dl.
(4) Dl, D2, and D3 are overlain by C. 41
Tt-e, S3 / t
D ll D2 Tt S2 S3, S2 tl ♦l e H It SI R3, R2
D3 S3 4
(a) (b)
Tt-e . S3
9
Fig. 24. Results of Markov chain analysis.
(a) "Standard" lithofacies sequence of the Terraba Formation, based on matrix values in Table A-4.
Heavy lines connect lithofacies transitions that occur much more commonly than those would characterize a random association (matrix value > 0.3). Solid lines connect transitions occur more commonly than those would characterize a random association (0.3 > matrix value > 0.1). Dashed lines connect transitions slightly more commonly than those wDuld characterize a random association (0.1 > matrix value > 0.035).
The one-way arrows represent the preferred relation between adjacent lithofacies. For example, D3 overlies C; C overlies Dl, D2, and D3.
The two-way arrows indicate the relation between interbedded lithofacies. For example, Lithofacies F is more likely to underlie D2 and D3 than to overlie them because solid arrows point away from F into D2 and D3, while dashed arrows point to F from D2 and D3.
(b) Transition of "sediment gravity flow divisions" based on corre sponding lithofacies in (a). See Table 4 for detail. 42
Table 5. Transition between sediment gravity flow divisions as sug gested by Figure 24. Within each row, transition of divisions from right to left (R2 to Te) represents downslope changes in turbidite organization (Lowe, 1982). It also represents an idealized verti cal transition of divisions deposited from turbidity currents. Vertical changes of lithofacies at left (E-'-Dl^C) are derived from left and central part of Figure 24a. Vertical changes of litho facies at right (C-*-D3-*“F*-D2*-C) are derived from central and right part of Figure 24a.
Litho- Corresponding sediment Litho- Corresponding sediment facies flow divisions facies flow divisions
C Te Td Tt S3 S2 SI R3 R2 Te Td Tt S3 S2 SI R3 R2 4
Te Td |Tt S3 S2 SI R3 R2 D2 Te Td Tt S3 S2 SI R3 R2
Te Td Tt S3 S2 SI R3 R2 Te Td Tt S3 S2 Si R3 R2
t>3 Te Td Tt S3 S2 SI R3 R2 4
Te Td Tt S3 S2 SI R3 R2
These results indicate that lithofacies transitions from high- density turbidite divisions (R2, R3, Si, etc.) to low-density divisions
(Tt, Td, Te, etc.) occur more commonly than those from low-density divi sions to high-density divisions. This preferred transition from high- to low-density divisions is expected because tubidites are deposited partly as a consequence of waning current velocity, and high-density de position normally precedes low-density deposition (Lowe, 1982). For example, if l-*-m— »-n represents a vertical or lateral sequence deposited from a turbidity current, where 1 represents the initial deposit and n represents a later-stage deposit, then apparently in a composite se quence l-^nr-^n-^l-»-m-*-n-4'l--*“m-^'n there are more l-»-m and m-*~n (6 total) than n— ►! transitions (2 total). This also explains why there are more 43
E—►Dl, F—1HD2, and F-HD3 than E-»—Dl, F*"—D2, and F-®—D3, respectively
(Table 5), because transitions in the former represent sequential depo sition from individual waning currents, whereas the latter represents transitions between separate currents. One important implication of these results (Table 5) is that sedimentation in the study area was characterized by deposition from currents of progressively declining velocity, i.e., by tubidity currents, separated by long intervals of quiet, low-energy deposition from suspension.
The above interpretations are contingent on two important assumptions:
(1) There are no major erosional unconformities within the Terraba
Formation.
The vertical succession from one lithofacies to another
suggests that the two lithofacies represent once laterally adjacent
environments (Walther's Law), or, in the case of turbidite deposi
tion, adjacent lithofacies are genetically related. The presence of
an erosional break may therefore lead to biased interpretations
(Middleton, 1973).
In the study area, deposition of Lithofacies E (cross-bedded
sandstone) and F (conglomerate), both having erosional lower con
tacts, was commonly preceded by erosion of the underlying sediments,
but this erosion was probably limited to areas within channels and
did not remove significant thicknesses of strata. The occurrences
of high-energy lithofacies (E and F) also are few (4% of measured
sections), indicating that erosional breaks, if present, are negli
gible. In contrast, perhaps very little erosion was associated with
the deposition of the other two lithofacies (C and D). (2) The four clastic lithofacies of all measured sections are
genetically related.
If the four lithofacies were deposited in unrelated environ
ments, then obviously the statistically compiled lithofacies se
quence (Fig. 24) cannot be applied to environmental interpretation.
Field work shows that three of the four lithofacies occur in one
measured section (PY-57), and that all four lithofacies occur in
closely-spaced sections 66, 67, 68, 69, 70, 71, and 72, indicating
that the four Terraba lithofacies are closely associated at least in
the above sections.
In order to further test the genetic relationship among the
sandstone layers, a cumulative thickness distribution curve of all
sandstone beds from the Terraba Formation is presented in Figure 25.
The curve shows close approximation to a log-normal distribution,
suggesting that all sandstone beds belong to the same layer
thickness population. Similar reasoning has been used by Corbett
(1972), Moore (1973), and Hiscott and Middleton (1979) to suggest a
common process for the deposition of deep-water sandstones. It is
therefore reasonable to apply the lithofacies sequence (Fig. 24) to
exposed sections in the whole study area, including places where
only some of the four lithofacies are present. 45
_j_____ i____ I____ I____ I____I .J____ L O 1 2 3 4 5 6 7
THICKNESS = 2* • 2.5 cm
Fig. 25. Cumulative thickness distribution for all sandstone layers of the Terraba Formation, based on 198 measurements of layer thickness and 473 m of measured sections. SEDI. ENTARY PETROLOGY
Studies of the petrology of recent sediments and ancient sediment ary rocks indicate that each major tectonic setting is associated with a compositionally distinctive suite of sediments (Dickinson, 1974;
Dickinson and Suczek, 1979; Dickinson and Valloni, 1980; Valloni and
Maynard, 1981; Maynard et al., 1982). Furthermore, the evolution of an individual tectonic terrane can commonly be reconstructed by studying the compositional evolution of sediments deposited in adjacent basins
(Dickinson and Rich, 1972; Ingersoll, 1978).
Detrital modes of sandstones in the study area were studied using petrographic techniques outlined by Dickinson (1970), Dickinson and Rich
(1972), and Graham et al. (1976). Thin sections were prepared from 44 samples of sandstones of the Terraba Formation and 300 framework grains were counted in each. Because volcanic rock fragments make up more than
60% of the framework grains, they were further subdivided into three categories (Dickinson, 1970): vitric (v), microlitic (m), and lathwork
(1).
In addition, selected plagioclase and pyroxene grains were individ ually analyzed using electron microprobe techniques. Plagioclase compo sition is useful in determining the fractionation trends in magmas
(Ewart, 1976; Gill, 1981). Analysis of pyroxene can provide information on petrogenetic evolution, which is commonly the only primary mineral that still preserves such information (Garcia, 1978).
46 47
Petrographic Study
The point count results are presented in Table A-5 and recalculated
detrital modes in Table A-6. The Q-F-L plot (Fig. 26) indicates that
sandstones in the study area are composed almost exclusively of volcanic
rock fragments and plagioclase feldspar grains. Quartz and potassium
feldspar are rare to absent. Quartz is present in only 8 of 44 samples
and makes less than 0.7% of all samples where it is present except one.
A single sample (# 70-1) was found containing 16% quartz. It occurs
only as volcanic quartz characterized by embayed, non-undulose grains with straight boundaries intersecting at 60 or 120 degrees. The sand
stones are commonly cemented by authigenic calcite and/or zeolites and most have a detrital matrix content below 5%. They can be classified as
lithic graywacke to subgraywacke (Pettijohn, 1957) or feldspathic
litharenite to volcanic arenite (Folk, 1974).
Q-F-L plots (Fig. 26) of the sandstone samples indicate that the
sediments were derived directly from a magmatic arc provenance
(Dickinson and Suczek, 1979). The predominance of plagioclase and volcanic rock fragments, the lack of plutonic and metamorphic debris, and the absence of sedimentary rock fragments indicate that sediments in the study area were derived from a first-cycle undissected volcanic arc with no observable influence from continental rocks. The extremely immature texture and composition indicate no extensive physical
(transportation) or chemical (weathering) breakdown prior to deposition.
Sandstones in the study area were deposited over a relatively short portion of the Tertiary Period (latest Oligocene to early Middle
Miocene). Finer subdivision of the litho-stratigraphic units and 48
Q
RECYCLED OROGEN
PROVENANCE
MAGMATIC ARC
PROVENANCE
mean
F 4— ^
Fig. 26. Q-F-L diagram showing the composition of sandstone samples in the study area based on point count data of 44 thin sections.
Each dot represents one thin section. Number of tick(s) on some dots indicating the number of additional thin section(s) with the same Q-F-L values. Mean value is indicated by arrow.
Enclosed areas showing mean framework modes for sandstone suites from different types of provenances (Dickinson and Suczek, 1979).
Q: Total quartz grains. F: Total feldspar grains. L: Total lithic fragments. 49
correlation between sections is beyond the resolution of available bio-
stratigraphic data. Poor exposures and possible structural complexities
also prevent direct physical correlation. It was not possible, there
fore, to attempt a statistical search for systematic trends in sandstone
composition. In their studies of detrital modes of sandstone samples at
central and western Terraba Trough, Hutchison (1980) and Phillips (1983)
recognized a slight upward increase in the proportion of vitric volcanic
rock fragments (v'), and a corresponding decrease in the abundance of
plagioclase grains (P). Their samples, however, range in age from Late
Eocene to Late Miocene. The bulk of the vitric fragment-rich and
plagioclase-poor sandstones reported in these studies occurs in the
Curre Formation which is not exposed in the study area.
Although the temporal or spatial variation, if any, of sediments in
the study area can not be established due to uncertainties in correla
tion between sections, the overall lack of quartz and potassium feldspar
and the high percentage of vitric volcanic rock fragments indicate that
the sediments were derived from a compositionally rather homogeneous magmatic arc characterized by an explosive eruption style (Williams and
McBirney, 1979; Sigurdsson, 1982). There are no clearly discernable
sandstone samples or sections that do not fall within the overall
volcaniclastic sandstone population. The presence of some volcanic quartz- and plagioclase-rich units appears to reflect local sources
rather than fundamentally different source terranes. Microprobe Analysis
In order to more accurately determine the composition of the source
rocks providing sediment to the study area, plagioclase and pyroxene
\ grains in the sandstone samples were analyzed by using electron micro
probe. Twenty four samples were chosen for analysis; 22 of them are
from outcrops of the Terraba Formation in the study area, one from the
Quaternary Brujo Formation east of San Vito (location 37) and one from
Miocene sandstone (Terry, 1956) north of David, Panama, which represents
the eastern extremity of the Terraba Trough.
Analysis of 148 plagioclase grains shows that the anorthite content
ranges from An28 to An88 (Fig. 27). The predominance of andesine and
labradorite indicates that the source rock was of basaltic andesite to
basalt composition (Ewart, 1976, 1982; Gill, 1981).
Consideration of plagioclase composition of sandstone samples
versus their distribution in the study area (Fig. 28) shows no apparent
trend of magma fractionation. Assuming that the composition of the
sampled volcaniclastic rocks is representative of the composition of the
coeval eruptives, then the lack of systematic variation of anorthite
content suggests that the composition of the magma remained essentially
unchanged. The same conclusion was reached by Hutchison (1980) and
Phillips (1983) in their analyses of samples from the central and western Terraba Trough.
Analytical results of pyroxene samples (Fig. 29) show that most
pyroxenes fall within the field of augite and that they range within
Wo33-48 En38-50 Fs7-23. This corresponds to values for augite pheno-
crysts in andesite (Gill, 1981). Figure 29 indicates that pyroxene 51
ANDESINE 60 S'*
40 LABRADORITE £ Frequency% n = 148 30
20 “
10 BYTOWNITE OLIGO- CLASE
f ...... i 0 10 30 50 70 00 100
Anorthite Content (Mol %)
Fig. 27. Anorthite content of detrital plagioclase from sandstones of the study area based on electron microprobe analysis.
Core or rim values are not specified because zoning is not evident in these sand-sized plagioclase grains.
Standards used for comparison: Lake County Plagioclase (Oregon) - USGS Microcline Standard - Australian National University COSTA RICA
STUDY AREA
-8"50 SAN VITO
-8*45' AGUA BUENA
166 60
-8*40'
INTERAMERtCAN
Km 82*55'
Fig. 28. Map showing maximum anorthite content (mol %) of detrital plagioclase from Terraba sandstones at representative locations.
Also shown are four locations ( X ’s) containing detrital pyroxene with alkaline affinities. Alkalic DIOPSIDI 8ALITE FERROSAL1TE
FERROAUGITE AUQITE • •
ENDIOPSIDE Calc-alkaline
Fig. 29. Pyroxene quadrilateral showing that most pyroxene samples are augite and have calc-alkaline trend.
Magma trends (Garcia, 1978, 1979) showing iron enrichment for tholeiitic magmas, increasing alkalinity for alkaline magmas, and only slight iron enrichment for calc-alkaline magmas.
Standards used for comparison: Kakanui Augite (New Zealand) - Smithonian collection Johnstown Hypersthene - Smithonian collection samples in the study area are plotted predominantly in the calc-alkaline field, with a few showing an alkaline trend.
A plot of Ti02 versus A1203 content for pyroxene grains (Fig. 30) shows that most pyroxene grains have a calc-alkaline affinity except five samples which lie in the alkaline field. Of these five samples, one is from the Quaternary Brujo Formation, one from Miocene rocks in
Panama, and three from the Terraba Formation which is latest Oligocene to early Middle Miocene in age in the study area (locations of the four
Costa Rican alkaline samples are shown in Fig. 28). This indicates that alkaline volcanism started early in this part of the Terraba Trough; elsewhere in Costa Rica alkaline rocks did not appear until Pliocene and
Pleistocene time (Pichler andWeyl, 1975; Hutchison, 1980; Weyl, 1980;
Lowery, 1982; Phillips, 1983). Fig. 30. Plot of titanium versus aluminum for detrital pyroxene grains from the study area, in comparison with those from central and western Terraba Trough by Hutchison (1980), Lowery (1982), and Phillips (1983).
Curve separates alkaline from nonalkaline source (after Garcia, 1978, 1979). Five samples are in the alkaline field. 56
Paso Real Fm./Brujo Fm • Terraba Fm./Curre Fm,
Alkaline
THIS STUDY
Non-alkaline
2.0 4.0 6.0 8.0 10.0
Alkaline
. HUTCHISON 1980
Non-alkaline T i0 2
W t . % 2.0 4.0 6.0 8.0 10.0
Alkaline
LOWERY 1982
Non-aikaline 4.0 10.0
. PHILLIPS 1983
••
Non-alkaline
2.0 40 6.0 8.0 10.0 ai2o 3 wt.% TECTONICS AND BASIN EVOLUTION
Basin Setting
Available on-shore data on the geology of Central America indicate
that the area of Terraba Trough has been located near an intra-oceanic
arc-trench system since at least Late Cretaceous time (Malfait and
Dinkelman, 1972; Dickinson and Coney, 1980; Pindell and Dewey, 1982;
Lundberg, 1983). Offshore studies further suggest that the southern
Central America arc-trench system developed in the Eocene due to the
breakup of the East Pacific Plate into the Farallon and Caribbean Plates and was subsequently characterized by subduction of the Farallon Plate beneath the Caribbean Plate (Hey, 1977; Lonsdale and Klitgord, 1978).
Sedimentary basins developed near active ocean-ocean plate margins include those located between the trench and the magmatic arc (fore-arc region basins), within arcs (intra-arc basins), and behind the arc
(back-arc basins) (Dickinson, 1974). Based on measurements of the width of arc-trench gap in 20 active modern convergent systems, Dickinson
(1973) showed that the spacing between arc and trench increases system atically during the life of a given convergent junction, provided that the polarity of subduction is maintained without change or significant interruption. The relative steadiness in polarity of the post-Eocene convergent system in the area of southern Central America (Hey, 1977;
Dickinson and Coney, 1980) implies that the Eocene-Oligocene arc was located between the Talamanca, site of arc plutonism between Late
57 Oligocene and Late Miocene time (Weyl, 1980), and the Nicoya Complex
which lies adjacent to the modern trench. If the Eocene-Oligocene arc
was coincident with or immediately in front of the modern Talamanca, the
Terraba Trough was then, as now, in a fore-arc setting. However, if the
Eocene-Oligocene arc was closer to the Nicoya Complex, the Terraba
Trough could represent either an intra- or a back-arc basin.
If the study area was situated in an intra-arc setting, volcanic
units, including both flows and pyroclastic deposits, associated
intrusive rocks, and remnant eroded cone aprons should occur at least
locally within the Terraba Trough sequence. With the exeption of minor
Late Miocene intrusives along the Rio Terraba northwest of the study
area and a single unit of Early Miocene flow breccia near Alto Garrote
(Fig. 3), Eocene-Oligocene igneous rocks and frontal arc facies
(Hutchison, 1980) do not occur in the Terraba Trough.
Weyl (1957) has described lower Tertiary olivine basalt alternating with tuffs in the Cordillera de Talamanca. Lower Tertiary volcanic rocks have also been reported from both the Pacific and Caribbean sides of the Talamanca by Dengo (1962, Fig. 4). In the Limon Basin, Eocene
tuffs, agglomerates, and flows are interbedded with Upper Eocene marine
sediments (Dengo, 1962). These occurrences have led De Boer to suggest
that the magmatic arc in Costa Rica has migrated away from the Pacific, with the Eocene arc located immediately in front of and on the Pacific
side of the Talamanca (De Boer, 1979, Fig. 3). If so, the Terraba
Trough has been in the fore-arc region since at least Eocene time.
Deposition in fore-arc region can take place in one of several basin settings: trench fill, slope basin, fore-arc basin, and inter massif basin (Dickinson and Seely, 1979; Dickinson, 1982). The general ly coarse grain size, large areal extent, and absence of intense thrust
ing and blueschist metamorphism of the sedimentary rocks in the Terraba
Trough suggest that the study area represents a fore-arc basin.
Geologic Development
Rocks exposed in the eastern Terraba Trough belong to the Brito,
Terraba, Paso Real (?), and Brujo Formations (Fig. 4). Other sedi
mentary units of southwestern Costa Rica, including the Nicoya Complex,
Curre and Paso Real Formations, and a Pliocene mudstone unit, although
not found in the study area, are well exposed in adjacent parts of the
Terraba Trough and are genetically inseparable from rocks in the study
area. Together they record the evolution of the Tertiary fore-arc
basin. The following discussion of the geologic history of this part of
Costa Rica is based on results obtained from both the study area and
other parts of the Terraba Trough. The history is discussed in five
intervals (Fig. 31): Late Jurassic to Early Eocene, Middle Eocene to
Late Eocene, Oligocene to early Middle Miocene, Middle to Late Miocene,
and Pliocene to Holocene.
I - Late Jurassic to Early Eocene: Development of Nicoya Complex
The oldest rocks in Costa Rica belong to the Upper Jurassic to
Upper Cretaceous Nicoya Complex. They are exposed discontinuously along
the Pacific coast with the largest outcrop areas on the Nicoya and Osa
Peninsulas. Similar rocks also crop out on the Pacific coast of Panama west of the Canal Zone. East of the Canal Zone (Fig. 32), Nicoya 60
\
Fig. 31. Correlation between major tectonic events and evolution of the Terraba Trough. Absolute time scale after Palmer (1983). RELATIVE IGNEOUS Relative WATER DEPTH GEOLOGIC EVENTS ACTIVITY SEDIMENTATION AGE convergence rate
(Ma) cm/yr SHALLOW DEEP LOW HIGH HOLOCENE Brujo Fm. Uplift of Talamanca 0. 01- Cocos Ridge clogged M. America PLEISTOCENE \m zm m Trench Formation of Terraba Trough Paso Real PLIOCENE Fm. 5.3- Coarse elastics in shallowing environments 11.2 MIOCENE 16.6- Terraba Fm. 11.2 Deep-water clarrtics 23.7 Farallon Plate broke 14) into Cocos & Rio C laro 7.5 Nazca Plates 30.0 OLIGOCENE Mbr. Eustatic sea level drop 6.2 Carbonates on local shoals 36 j asTOunded by deep water E. Pacific Plate broke 14) into 40.0 Brito Fm. EOCENE (1) Farallon & (2) Caribbean Plates 52.0 (1) subducted under ( 2 )
57.8
PALEOCENE
66.4 Early fragmentation of E. Pacific P late ? < « 1-3 Uplift & deformation of Nicoya u j o Nicoya Complex, andesite volcanism CC 111 00 Complex Pelagic to hemipelagic 144
208 62
1 5 0 Ion NICARAGUA
CARIBBEAN
COSTA RICA
NICOYA
»0. MA
OSA
PACIFIC COLOM- n/.: :\BIA
Fig. 32. Distribution of "Nicoya Complex" (dotted pattern) and Eocene shallow-marine limestone (X's and solid pattern) in Costa Rica, Panama, and Colombia. The distribution of limestone in Panama may not be as continuous as shown.
Source:
"Nicoya Complex" - Case, 1974; Schmidt-Effing, 1979 (modified); Weyl, 1980. Eocene limestone in Costa Rica - Cole, 1953; Malavassi, 1961; Henningsen, 1966a; Rivier, 1973; Lundberg, 1982; Gose, 1983; E. Malavassi, personal commun. Eocene limestone in Panama - Terry, 1956; Weyl, 1980. Complex-type rocks are found both on the Caribbean and Pacific coasts
(Case, 1974).
The Nicoya Complex is made up of tholeiitic basalt and interlayered radiolarite, hematitic claystone, and manganiferous units (De Boer,
1979; Galli-Oliver, 1979; Schmidt-Effing, 1979; Kuijpers, 1980;
Lundberg, 1982; Lew, 1983). The basalt is chemically similar to that erupted at spreading centers (Pichler and Weyl, 1975). The associated sediments and organisms are characteristic of the modern seafloor.
Thus, the Nicoya Complex probably represents uplifted oceanic crust
(Stibane et al., 1977; Schmidt-Effing, 1979).
The Nicoya Complex was deformed during the Santonian to Campanian
(middle Late Cretaceous)(De Boer, 1979; Kuijpers, 1980). An angular unconformity separates the Upper Jurassic to Upper Cretaceous basalt and associated sediments from overlying Campanian to Lower Eocene rocks composed of pelagic to hemipelagic sediments, volcaniclastic turbidites, and shallow-marine limestone. The Campanian also marks the first appearance of abundant andesite clasts in the sequence. The deformation of the Nicoya Complex, the presence of the angular unconformity, and the initiation of andesitic volcanism, all occurring from Santonian to
Campanian, suggest that the Nicoya Complex was at this time uplifted and located near a magmatic arc (Kuijpers, 1980). The development of the magmatic arc suggests that fragmentation of the East Pacific Plate and associated subduction began as early as in Late Cretaceous time
(Galli-Oliver, 1979; Pindell and Dewey, 1982).
A number of mechanisms have been proposed for the deformation and uplift of the Nicoya Complex. Stibane et al. (1977) suggested that descending oceanic lithosphere, upon approaching the subduction zone, was first vertically faulted and differentially uplifted in the area of the outer swell (morphological term after Karig and Sharman, 1975). The subduction zone then shifted offshore away from the arc. The faulted and differentially uplifted blocks, originally belonging to the descend ing lithosphere, were thus accreted to the overriding lithosphere.
Schmidt-Effing (1979) interpreted the Nicoya Complex as part of an oceanic plateau. Due to its low density, the plateau was obducted when it reached the subduction zone. Galli-Oliver (1979), Karig (1982b), and
Lundberg (1982) suggested that the Nicoya Complex represents an uplifted trench slope break or subduction complex composed of slabs of oceanic crust that were progressively sliced off of the descending oceanic lithosphere.
The contact between the Nicoya Complex and the Brito Formation is not exposed within the Terraba Trough. Consequently, the relationship between these terranes must be inferred indirectly. It is unclear, for instance, whether the Nicoya Complex has been recently accreted against the southwest side of the Terraba Trough or whether it has been in proximity over a much longer period of geologic time. Paleomagnetic studies of sedimentary rocks ranging in age from late Cretaceous to
Eocene collected from various sites in Costa Rica, including the Nicoya
Peninsula and Terraba Trough, show that late Mesozoic and early Tertary rocks throughout western Costa Rica have the same paleomagnetic pole
(Gose, 1983). This suggests that the relative latitudinal positions of the Nicoya Complex and rocks in the Terraba Trough was the same in
Eocene as it is now. This paleomagnetic constraint and the close association between the outcrops of Nicoya Complex and Eocene shallow- marine limestone (Fig. 32), which is in stratigraphic continuity in the Terraba Trough with younger Tertiary units, suggest that the Nicoya
Complex has been spatially associated with sediments of the Terraba
Trough since at least Eocene time.
II - Middle Eocene to Late Eocene: Initiation of Fore-arc Basin
Following Santonian-Campanian deformation and uplift of the
Nicoya Complex, a second phase of tectonic activity took place in Middle
Eocene (Krushensky et al., 1976; Schmidt-Effing, 1979; Kuijpers, 1980,
Lundberg, 1982). This is suggested by the existence of an angular unconformity on the Nicoya Peninsula separating folded and faulted post-
Campanian, pre-Eocene rocks from overlying largely undeformed Miocene rocks (Kuijpers, 1980). On the Nicoya Peninsula and over wide areas in southern Central America, regional shoaling at this time resulted in the deposition of Eocene shallow-marine limestone (Cole, 1953; Terry, 1956;
Malavassi, 1961; Henningsen, 1966a; Rivier, 1973).
The distribution of this limestone (Fig. 32) may provide important clues to the tectonic evolution of southern Central America. Inspection of this distribution suggests an association between the Nicoya Complex and Eocene limestone. In all cases where information is available, the presence of Brito limestone correlates closely with exposures of the
Nicoya Complex. There are no Nicoya Complex-type rocks exposed in
Nicaragua, and there the Brito Formation is composed predominantly of clastic rocks and a greatly subordinate amount of limestone (Weyl,
1980). Also, thick Eocene limestone occurs only in the central and eastern parts of the Terraba Trough where the Nicoya Complex is exposed on the nearby Osa Peninsula, whereas little limestone is present in the western Terraba Trough (Henningsen, 1966a; Hutchison, 1980; Phillips,
1983). In the latter area, Eocene rocks consist largely of volcani- clastic turbiditic units, some containing redeposited shallow-marine larger foraminifera (Phillips, 1983).
Occurrences of limestone in central Costa Rica (X's in Fig. 32) are exceptions; they are far inland from any exposed Nicoya Complex.
It has been reported that a diagonal belt across central Costa Rica, overlapping these "inland" occurrences of Eocene limestone, is geologi cally discontinuous from the rest of Costa Rica (De Boer, 1979). The relationships of these inland limestones is unknown; however, because their lower contact is apparently not exposed, they may also be under lain by rocks of the Nicoya Complex.
This close association suggests that the presence of Nicoya Complex rocks may have controlled deposition of the Brito limestone. The timing of the Eocene shoaling coincides with proposed fragmentation of the East
Pacific Plate in Middle and Late Eocene (Malfait and Dinkelman, 1972;
Dickinson and Coney, 1980; Sykes et al., 1982). This fragmentation occurred when the northward motion of the East Pacific Plate with respect to North American Plate was terminated by the collision of the
Greater Antilles island arc at the northern front of the East Pacific
Plate with the Bahama carbonate platform on the North American Plate
(Dickinson and Coney, 1980). As a result, the East Pacific Plate fragmented into the Caribbean and Farallon Plates, with the Farallon
Plate undergoing northward subduction beneath the Caribbean Plate
(Malfait and Dinkelman, 1972; Dickinson and Coney, 1980; Sykes et al.,
1982). It seems likely that rocks of the Nicoya Complex, which were previously uplifted in the Late Cretaceous, were again uplifted in the 67
Eocene as a result of this fragmentation and the initiation of
subduction. The uplifted blocks of Nicoya Complex apparently served as
shallow-water ridges on which organisms flourished and carbonates
accumulated. In intervening areas, such as the western Terraba Trough,
clastic sequences were deposited. That some volcanism accompanied the
subduction is indicated by the presence of Eocene volcaniclastic units
both locally within the Brito limestone (Henningsen, 1966a) and as a
separate, largely volcaniclastic facies (Phillips, 1983). The overall
sedimentation in Middle to Late Eocene was probably characterized by
shallow-water carbonate deposition associated with locally uplifted blocks of the Nicoya Complex, and by deeper water deposition in areas between the uplifted basement blocks (Fig. 31).
Ill - Oligocene to Early Middle Miocene : Subsidence of Fore-arc Basin and Alkaline Volcanism
Lower Oligocene rocks are not exposed in the study area. Instead, shallow-marine Eocene carbonates of the Brito Formation are overlain directly by Upper Oligocene shallow-water limestone of the Rio Claro
Member of the Terraba Formation. The absence of Lower Oligocene units may reflect exposure and erosion at about 30 Ma, near the middle of the
Oligocene, when global sea level may have dropped nearly 340 m (Vail et al., 1977). The Oligocene low-stand of sea level, although suggested mainly by evidence obtained from passive margin sequences, should have affected all shallow-marine areas, including the Terraba Trough, regard less of their tectonic setting. For example, stratigraphy of San
Joaquin Basin, California, which was located near an active margin, also reflects this Oligocene event (Vail et al., 1977). However, within the
Terraba Trough, direct evidence indicating subaerial exposure, such as
soil horizons or subaerial dissolution surfaces in the limestone, was
not found. The only reported Lower Oligocene rocks in southwestern
Costa Rica occur near the southwestern margin of the Terraba Trough in a
sequence that is structurally below rocks of the Middle-Upper Eocene
Brito Formation (Henningsen, 1966a). It is composed of claystone and
interbedded fine-grained, thin bedded turbidites and has been telescoped
and emplaced by northeasterly-dipping thrust faults (Mora, 1979). The
sedimentological relationship of this sequence to others in the Terraba
Trough is unclear. If the emplacement occurred in post-Oligocene time,
underthrusting has moved rocks from more southwesterly areas into posi
tions beneath those of the more proximal Terraba Trough sequence; then
these Lower Oligocene rocks may represent distal units deposited in
deeper water nearer to the trench. A sliver of Lower Miocene rocks
occurring near Villa Neily (Fig. 4) probably has the same origin.
When deposition resinned after the Lower Oligocene hiatus, shallow- water carbonate rocks of the Rio Claro Member were formed within the
study area. Elsewhere, rocks of Late Oligocene age consist mainly of dark, pyrititic claystone and thin turbiditic sandstone (Henningsen,
1966a; Phillips, 1983). The overall Late Oligocene depositional pattern appears to be similar to that prevailing during the deposition of the
Eocene Brito Formation which was characterized by carbonate sedimenta tion on local highs underlain by rocks of the Nicoya Complex and clastic deposition in laterally associated deep-water areas.
Deposition of the Upper Oligocene Rio Claro Member was followed by rapid subsidence in the latest Oligocene. In the measured section PY-75, Rio Claro Member limestone is overlain directly by Upper Oligo cene mudstone containing deep-marine benthic foraminifera Melonis pompilioides. From late Late Oligocene to early Middle Miocene, the area was a deep-marine trough characterized by turbidite sedimentation.
A number of mechanisms have been proposed to account for subsidence or uplift in fore-arc regions:
(1) Thermal contraction of the overriding lithosphere associated with
the initiation of subduction or an increase in the rate of subduc
tion. Both of these mechanisms provide efficient means of cooling
the mantle below the overriding lithosphere, resulting in subsidence
(Hsui and Toksoz, 1979; Langseth et al., 1981). Langseth et al.
(1981) proposed that the late Paleogene subsidence of the Japan
trench/fore-arc region of northern Honshu was the result of initia
tion of subduction in its vicinity.
(2) Gravitational loading by sediment and water (Walcott, 1972).
Gravitational loading probably accounts for only a small amount
of subsidence in the Terraba Trough after the deposition of the
shallow-marine limestone because the subsidence rate outpaced the
accumulation rate of the sediments.
(3) Subduction erosion of the overriding lithosphere by Jthe descending
lithosphere. Erosion of the upper lithosphere can occur either at
the leading edge of the overriding plate, resulting in a truncated
margin; or along the base of the overriding lithosphere, resulting
in subsidence (Karig, 1974; Scholl et al., 1980; Langseth et al.,
1981). The amount of subsidence in the Terraba Trough caused by
subduction erosion can not be calculated at this time.
(4) The subduction of anomalous, low-density lithosphere, including aseismic ridges, intraplate island-seamount chains, or very young
lithosphere. Displacement of the asthenospheric wedge by this
anomalously low-density lithosphere may cause subsidence (subcrustal
loading) (Cross and Pilger, 1978, 1982; Pilger, 1981). This mecha
nism, however, will normally cause subsidence in the back-arc area
and uplift in the fore-arc region (R. H. Pilger, personal commun.),
and therefore was probably not effective during the early evolution
of the Terraba Trough.
(5) Underplating. This process, called "subcretion" by Karig and Kay
(1981), requiring the detachment of sediment from the descending
lithosphere and its accretion to the bottom of the accretionary
complex, has been used to explain the uplift without significant
deformation, of many mid- and upper-slope areas (von Huene, 1984).
However, indirect evidence for underplating has proved more convinc
ing than direct evidence (Karig, 1982a) and the effectiveness of
this mechanism to cause uplift is unknown.
(6) Subduction of lithosphere of different ages. The average density of
lithosphere is a function of age. Younger lithosphere is relatively
more buoyant than older, denser lithosphere (Molnar and Atwater,
1978). The fore-arc region may experience differential uplift or
subsidence due to contrast in buoyancy between different parts of
the descending lithosphere. Fore-arc region above the older litho
sphere may undergo subsidence relative to the neighboring fore-arc
region overlying younger lithosphere.
Of the above processes affecting the subsidence of the Terraba
Trough, (2) is unimportant, (3) is possible but speculative, (4) and 71
(5) are probably Ineffective in causing subsidence in fore-arc regions,
and there is little evidence for major differential uplift or subsidence
in different parts of the Terraba Trough (6). The most probable
mechanism for causing rapid Late Oligocene subsidence of the Terraba
Trough area and formation of the deep-water fore-arc basin was, there
fore, the initiation of subduction or an increase in the rate of
subduction.
From Middle Eocene to Late Oligocene, the Terraba Trough area was
characterized by the presence of local shoals and closely juxtaposed
deeper water environments. Not until the latest Oligocene did a single,
large, deep-water basin developed. Since the breakup of the Pacific
Plate into the Farallon and Caribbean Plates and the initiation of
subduction occurred much earlier, in the Eocene, the Late Oligocene
rapid subsidence of the fore-arc basin was most probably related to an
change in the rate of subduction. The Middle Eocene to Late Oligocene
period of limited subsidence, shallow-water carbonate sedimentation, and
subdued volcanic activity most likely reflect corresponding relatively
lower subduction rates.
Within the Upper Oligocene to lower Middle Miocene Terraba
Formation, in most sections, the thickness and/or number of sandstone beds as well as the grain size, amount of carbonized wood fragments,
Turritella, and Pecten appear to increase up section. These features suggest that the basin was filled and that the nearshore area which served as a sediment source of turbidite debris prograded towards the study area. In central and western Terraba trough, Henningsen
(1966a) also noticed that the overall grain size and thickness of sand stone beds of the lower part of the Terraba Formation (Oligocene) is 72
significantly finer and thinner, respectively, than those of the upper
part of the unit (Lower to lower Middle Miocene).
The rapid subsidence and development of a large, coherent fore-arc
basin in the latest Oligocene, and the upward change in lithology of the
Terraba Formation may be related to the major reorientation of plate movements in the area. At approximately 26 Ma (Handschumacher, 1976) to
29 Ma (Atwater, 1970), a segment of the Pacific-Farallon ridge collided with western North America. Consequently, at 25 Ma (Hey, 1977) to 27 Ma
(Lonsdale and Klitgord, 1978) the Farallon Plate broke up into the Cocos and Nazca Plates separated by the newly formed Galapagos Rift Zone
(Wortel and Coetingh, 1981; Schilt and Karig, 1982). The relative motion of the Cocos Plate to the Caribbean Plate became northeastward versus the east northeastward direction of the Farallon Plate prior to its breakup (Hey, 1977). Subduction rate increased from 7.5 cm/yr in
Late Oligocene to 11.2 cm/yr in Early Miocene as a result of this change from a oblique to a more normal subduction (R. H. Pilger, unpub. data).
These data support the inference that a rather drastic increase in the rate of subduction during the latest Oligocene may have triggered rapid subsidence in the Terraba Trough. An accompanying increase in the intensity of volcanic activity in the area due to the rising subduction rate may also have resulted in the introduction of more and coarser volcaniclastic debris into the basin.
During deposition of the Terraba Formation, there was not only an increase in the amount of volcaniclastic detritus but also the appear ance of a minor component of alkaline rocks. Results of the microprobe analysis of detrital pyroxene grains show that, in the study area, 4 of the 41 samples of Terraba Formation sandstone lie in the alkaline field 73
(Fig. 30). Furthermore, the four alkaline samples are located in widely scattered areas in the eastern Terraba Trough: two from Alto Garrote, one from north of Villa Neily, and one from Panama (Fig. 28). The general trend of magma evolution in Costa Rica from calc-alkaline to alkaline (Pichler and Weyl, 1975) is typical of many island arcs (Jakes and White, 1972; Miyashiro, 1974; Hawkins et al., 1984). However, among
198 pre-Pliocene samples analyzed from the central and western Terraba
Trough (Hutchison, 1980; Lowery, 1982; Phillips, 1983), none of them show alkaline affinity. The earlier appearance of alkaline magma in the study area and western Panama than in correlative rocks to the northwest calls for special explanation.
Two mechanisms have been suggested to generate alkaline magmas.
(1) DeLong et al. (1975) noticed that alkaline volcanoes occur (a)
near hinge faults at convergent boundaries, e.g., in the areas of
Bering, Grenada, Hispaniola, and Fiji; or (b) where a fracture zone
or other linear feature approaches the subduction zone with high
angle such as at Kanaga and Malaita. The above tectonic environ
ments created access to the sublithospheric region, normally not
tapped in island arcs, for extraction of alkaline magma.
(2) Sugisaki (1976) suggested that the lower the subduction rate, the
less the amount of water would be provided to the mantle by the
descending plate, and less heat, generated by friction between the
converging plates, would be produced to generate magma. The de
scending plate will therefore have to reach deeper levels to attain
the temperature and pressure required for melting. Under these
conditions of high pressure and nearly anhydrous melting, the magma
generated is alkaline. The progressive enrichment of alkalis with increasing depth to the Benioff zone was also reported by Dickinson
and Hatherton (1967), and was ascribed to fractionation at deeper
levels as well as to lower degree of partial melting.
Reconstruction of plate movements in the area of Central America
and Caribbean indicates that the East Pacific Plate fragmented into the
Farallon and Caribbean Plates in the Eocene time (Malfait and Dinkelman,
1972; Dickinson and Coney, 1980). The boundary between the Farallon and
Caribbean Plates, however, probably did not evolve into a subduction
zone everywhere simultaneously. From 48 Ma (Sykes et al., 1982) to 30
Ma (Lonsdale and Klitgord, 1978), their boundary near southern Central
America, for example, was a strike-slip fault. Nor was the direction of
subduction the same along the entire convergent boundary. Due to the
southwestward convexity of the convergent boundary, subduction along
southern Central America was probably highly oblique (Wadge and Burke,
1983). It seems likely that the study area, lying further south and
east than the central and western Terraba Trough, was in a region of
highly oblique subduction and may have been located close to a hinge
fault developed at a trench-transform fault junction. Both settings
would have favored the formation of alkaline magma. At the same time,
calc-alkaline eruptions characterized areas further to the northwest. IV - Middle to Late Miocene: Large-scale Igneous Activity and Basin Shoaling
Rocks of Middle to Upper Miocene Curre Formation, exposed in the
central and western Terraba Trough, are presumably buried beneath the
Pliocene Paso Real (?) and Quaternary Brujo Formations to the north and
northeast of the study area (Fig. 4). In the central and western
Terraba Trough, studies by Hutchison (1980), Lowery (1982), Phillips
(1983), and D. R. Lowe and R. H. Kesel (personal commun.) have shown
that the basin was filled rapidly during the Middle to Late Miocene.
This is indicated by vertical changes in the 830 m-thick Curre Formation
from deep-sea fan facies near the base to shallow-marine, fluvial, and
lacustrine facies near the top (Lowery, 1982). Throughout Costa Rica,
this interval of the Tertiary was characterized by intensive volcanism
and associated plutonism. It was during this time that the bulk of the
Talamanca intrusive rocks were emplaced (Dengo, 1962), and Lowery (1982)
has documented an abundant influx of fresh volcaniclastic and pyro-
clastic debris into the western portion of the Terraba Trough. The
increased level of magmatic activity and correspondingly high rate of
clastic sedimentation in the Middle and Late Miocene may be related to an increase in the relative convergence rate between the Cocos and
Caribbean Plates from 11.2 cm/yr in Early Miocene to 12.0 cm/yr in Late
Miocene time (R. H. Pilger, unpub. data). The direct correlation between high convergence rates and extensive arc magmatism is well documented (Donnelly, 1973, 1975; Kennett and Thunell, 1975; Kennett et al., 1977) and offers perhaps the best explanation for Middle-Late
Miocene activity in the Terraba Trough region. At the same time, uplift and deformation accompanied with active volcanism occurred in Panama
(Terry, 1956).
V - Pliocene to Holocene: Regional Emergence
The Miocene magmatism and shoaling in the Terraba Trough was followed by a brief transgression in the early Pliocene. The resulting deposits consist mainly of fine-grained claystone and document an interval of subsidence and little volcanism. Throughout the remainder of the Pliocene and during the Quaternary, the Terraba Trough has been emergent and characterized by alkaline volcanism (Paso Real Formation) and alluvial and lacustrine environments (Paso Real and Brujo Forma tions) (Phillips, 1983; D. R. Lowe and R. H. Kesel, personal commun.).
In the deep-sea planktic foraminiferal record, the stratigraphic distribution of Globorotalia miocenica and Globorotalia multicamerata and the coiling trends of Neogloboquadrina pachyderma and Pulleniatina indicate that the seaway in southern Central America connecting the
Pacific and Caribbean was closed in the Pliocene after about 3.6 Ma
(Kaneps, 1970; Keigwin, 1978). SUMMARY AND CONCLUSIONS
The Tertiary sequence in Terraba Trough of southwestern Costa Rica includes more than 4,500 m of strata ranging in age from Middle Eocene to Holocene. The sequence represents a major transgressive-regressive cycle. From Middle Eocene to Late Oligocene time, deposition took place predominately in shallow-marine environment. In the Late Oligocene, the
Terraba Trough area subsided rapidly and was characterized by deep-water sedimentation. From Oligocene to Holocene, sedimentation occurred in sequentially shallower water depth, from deep- through shallow-marine, to fluvial, and lacustrine environments.
In the study area, the Eocene Brito Formation is 1,000 m-thick and made up of massive, white to yellowish gray, fossiliferous limestone.
Major allochems include coralline algae, larger benthic foraminifera, debris of echinoids, bivalves, and bryozoans, and small amount of planktic foraminifera. The light color, texture, and the fossil assemblage indicate that the Brito Formation was deposited on a shallow carbonate shelf in well-oxygenated water.
The Brito Formation is overlain directly by a thin carbonate unit of Late Oligocene age, here defined as the Rio Claro Member of the
Terraba Formation. It consists of two lithofacies : A and overlying
Lithofacies B. Lithofacies A is composed of rhodoliths in a matrix of bioclasts. The laminar growth form of coralline algae, the spheroidal shape, and the moderate to good sorting of the rhodoliths suggest that they were formed on a carbonate shelf swept periodically by storm currents. Lithofacies B consists of algal-foraminiferal grainstone and
77 is characterized by low-angle, tabular cross-sets. The characteristics
of the cross-sets, the large grain size of the allochems, and the
associated bioclasts indicate that Lithofacies B was deposited in
shallow marine with a water depth of at least 6 meters.
After the deposition of the Rio Claro Member, the study area
underwent rapid subsidence and was dominated by clastic sedimentation.
The resulting sediments make up the remainder of the Terraba Formation.
These clastic rocks were subdivided into four lithofacies, each charac
terized by distinctive lithology and association of sedimentary struc
tures. Results of Markov chain analysis of vertical relationships
between the lithofacies indicate that sedimentation in the study area
was characterized by deposition from turbidity currents, separated by
long intervals of quiet, low-energy deposition from suspension or
dilute, muddy turbidity currents.
The results of this study, combined with those from other parts of
the Terraba Trough, show that development of the fore-arc basin occurred
in three stages. (1) The absence of Lower Oligocene rocks and the
general characteristics of the Brito Formation and Rio Claro Member of
the Terraba Formation document an initial stage from the Middle Eocene
to latest Oligocene during which there was limited subsidence and that
at least parts of the area remained in a shallow-marine environment.
(2) This was followed in the latest Oligocene by a period of rapid sub
sidence during which the basin evolved into a deep trough. The trough was supplied with varying amounts of coarse, volcaniclastic detritus.
Terrigenous clastic strata of the Terraba Formation are characterized by
a vertical increase in sediment size and abundance of sandstone. These
trends appear to document a gradual shoaling of the basin from Late Oligocene to Middle and Late Miocene time. (3) The third stage of
Terraba Trough evolution, during the Pliocene and Quaternary, was one of uplift and the wide development of alluvial and lacustrine systems.
One of the most significant results of this study is the apparent close correlation between roagmatism and sedimentological evolution of the Terraba Trough and the inferred rates of convergence of lithospheric plates. Limited subsidence and mild volcanism during most part of
Eocene and Oligocene correspond to a low convergence rate between the
Farallon and Caribbean Plates. Rapid subsidence in the latest Oligocene and subsequent intense magmatism in Miocene time reflect the accelerated subduction of Cocos Plate beneath the Caribbean Plate.
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Scholl, D. W., Von Huene, R., Vallier, T. L. , and Howell, D. G., 1980, Sedimentary masses and concepts about tectonic processes at under thrust ocean margins: Geology, v. 8, p. 564-568.
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STUDY AREA
-8*50 SAN VITO
J. 73-2 j° ]73-3Ai
35-6
J \ 6 3 \ 2 - 8*45 75f 774-2^"-4' V 7 4 - 4 K AGUA BUENA
34-5P ,.31-1 <32-3 32-4
,56-4
Km 82*55
Fig. A-l. Map showing location of fossil samples listed in Appendix 1. Those not found in this map can be located in the described stratigraphic sections (Appendix 3). Refer to Figure 3 for location of described sections. APPENDIX 1 Identified Faunas in the Study Area
1. Brito Formation
29-8 31-1 32-3 31-2 32-4 75-4 34-5P
Discocyclina sp. X Eorupertia sp. X X Fabiania cubensis (CUSHMAN and BERMUDEZ) X X Heterostegina ocalana CUSHMAN X Lepidocyclina (L.) macdonaldi CUSHMAN X X X Lepidocyclina (]L. ) pustulosa DOUVILLE X X X Operculinoides trinitatensis (NUTTALL) X
Identifications by S. H. Frost (personal commun.)
Age: Middle to Late Eocene (S. H. Frost; E. Malavassi; B. K. Sen Gupta; W. A. van den Bold, personal commun.)
2. Claro Member (Terraba Formation)
74-1 74-2 63-2 63-4 75-1 75-5 75-2
Heterostegina antiIlea CUSHMAN X Lepidocyclina (E.) undosa CUSHMAN X Lepidocyclina (L.) mantelli (MORTON) X Lepidocyclina (L.) yurnagunensis CUSHMAN X X X Miogypsina sp. cf. panamensis (CUSHMAN) X X X X Pararotalia mexicana (NUTTALL) X X Pararotalia sp. cf. P. mexicana (NUTTALL) X Pararotalia sp. X Victoriella sp. X Carpenteria sp. X Astreopora goethalsi VAUGHAN X Porites panamensis VAUGHAN X
Identifications by S. H. Frost (personal commun.)
Age: Late Oligocene (S. H. Frost; E. Malavassi; B. K. Sen Gupta; W. A. van den Bold, personal commun.)
91 92
APPENDIX 1 (continued)
3. Terraba Formation
33-15 34-1 35-6 51-1 56-4
Catapsydrax stainforthi X BOLLI, LOEBLICH, TAPPAN Catapsydrax cf. dissimilis X (CUSHMAN and BERMUDEZ) Globigerina tripartita KOCH X Globigerinoides sicanus X DE STEFANI Globorotalia kugleri BOLLI X Globorotalia mayeri X X X X CUSHMAN and ELLISOR Globorotalia siakensis (LEROY) X X Globorotalia peripheroronda X X X X BLOW and BANNER Praeorbulina glomerosa (BLOW) X Praeorbulina transitoria (BLOW) X X
3. Terraba Formation (continued)
73-1 73-2 73-3 74-4 75-3
Globigerina angulisuturalis X X BOLLI Globigerina ciperoensis X X BOLLI Globigerinoides sicanus XXX DE STEFANI Globigerinoides trilobus X X (REUSS) Globigerinoides subquadratus XXX BRONNIMANN Globoquadrina altispira X (CUSHMAN and JARVIS) Globoquadrina cf. dehiscens X (CHAPMAN, PARR, COLLINS) Globoquadrina venezuelana X Globoquadrina sp. X Globorotalia mayeri XXX CUSHMAN and ELLISOR Globorotalia opima nana BOLLI X X Globorotalia opima opima BOLLI X Globorotalia cf. scitula X X (BRADY) 93
3. Terraba Formation (continued)
73-1 73-2 73-3 74-4 75-3
Globorotalia peripheroronda X X BLOW and BANNER Praeorbulina glomerosa (BLOW) X Allomorphina sp. X Bolivina sp. cf. 1$. mexicana CUSHMAN Bulimina pupoides D'ORBIGNY X Cassidulina sp. X X Cassigerinella venezuelana X Chiloguembelina sp. X Cibicides crebbsi X Melonis pompilioides X X X (FICHTEL and MOLL) Siphogenerina cf. transversa X Radiolaria X
Catapsydrax cf. dissimilis R (CUSHMAN and BERMUDEZ) Globorotalia opima opima BOLLI R R Globorotalia opima nana BOLLI R R Globigerina angulisuturalis BOLLI R R Globigerina ciperoensis BOLLI R R Melonis pompilioides R R Uvigerina sp. R R Chiloguembelina sp. R
R: Reworked species
Identifications by B. W. McNeely (personal commun.) Age (B. W. McNeely; W. A. van den Bold, personal commun.) 33-15: Early Miocene (N.7) 34-1: Early Miocene (N.7) 35-6: late Early to early Middle Miocene (N.7) 51-1: late Early to early Middle Miocene (N.7) 56-4: middle Early Miocene (N.6) 73-1: late Early Miocene to early Middle Miocene (N.7), with reworked early Late Oligocene species 73-2: late Early Miocene (N.7), with reworked early Late Oligocene species 73-3: late Early Miocene (N.7) 74-4: Late Oligocene (N.2 - N.3) 75-3: early Late Oligocene (N.2 - N.3)
Biostratigraphic zonation from Postuma (1971). APPENDIX 2
Markov Chain Analysis of Lithofacies in the Terraba Formation
Markov chain analysis is a statistical technique used for the detection of repetitive processes in space and time. Coal measure cyclothems and fluvial fining-upward cycles are good examples of litho facies successions which can be explained using Markovian processes.
This technique has been shown useful in stratigraphic analysis and in establishing facies models (Gingerich, 1969; Selley, 1970; Miall, 1973;
Cant and Walker, 1976; Hiscott and Middleton, 1979; Walker, 1979).
The first step in Markov chain analysis is to construct the "Litho facies transition count matrix" (Table A-l), in which the numbers of vertical transitions from one lithofacies to another are tabulated.
For example, number 7 at the intersection of row C (horizontal) and column D2 (vertical) represents that lithofacies D2 overlies lithofacies
C at 7 places in the measured sections. Lithofacies A and B (rhodolite and cross-stratified algal-foraminiferal grainstone) are not included in this analysis because their stratigraphic position is well known.
Table A-l. Lithofacies transition count matrix.
Lithofacies C D1 D2 D3 E F Row sum
C 0 3 7 11 2 1 24 D1 3 0 0 2 2 0 7 D2 7 0 0 4 0 2 13 D3 11 3 4 0 0 3 21 E 0 4 0 0 0 0 4 F 0 0 2 3 0 0 5
Total 74
94 95
Table A-l is then converted into a probability form (Table A-2).
The probability of D2 to overlie C is 7/24 =0.292 (Table A-2) because
lithofacies C is being overlain by different lithofacies at 24 places
(first row sum in Table A-l), 7 of which are by lithofacies D2.
In Table A-2, the sum of all probabilities in the same row equals to 1.
Table A-2. Lithofacies transition probability matrix.
Lithofacies C Dl D2 D3 E F Row sum
C 0 0.125 0.292 0.458 0.083 0. 042 1 D1 0.429 0 0 0.286 0.286 0 1 D2 0.538 0 0 0.308 0 0. 154 1 D3 0.524 0.143 0.190 0 0 0. 143 1 E 0 1.000 0 0 0 0 1 F 0 0 0.400 0.600 0 0 1
Table A-3 is then drawn up based on the assumption that all the
Terraba lithofacies occur in random order and are equally abundant.
Constructing Table A-3 is similar to placing all occurrences of Terraba
lithofacies into a box, drawing them out one at a time, and stacking
them in succession to form a stratigraphic column. The vertical succes
sion probability of one lithofacies succeeding another is then calcu lated. In this "Independent trials probability matrix" (Table A-3), the probability of lithofacies D2 overlying C, if they occur randomly, equals to "total number of lithofacies D2"/"total number of non-litho- facies C". Here the probability is 13/(74-24)=0.260. Table A-3. Independent trials probability matrix.
Lithofacies C Dl D2 D3 E F Row sum
C 0 0.140 0.260 0.420 0.080 0.100 1 Dl 0.358 0 0.194 0.313 0.060 0.075 1 D2 0.393 0.115 0 0.344 0.066 0.082 1 D3 0.453 0.132 0.245 0 0.075 0.094 1 E 0.343 0.100 0.186 0.300 0 0.071 1 F 0.348 0.101 0.188 0.304 0.058 0 1
A difference matrix (Table A-4) is calculated by subtracting Table
A-3 (random probability) from Table A-2 (observed probability). In
Table A-4, the positive values thus represent those lithofacies transi tions which have a higher-than-random probability of occurring, whereas the negative values represent lithofacies transitions much less common than random. For example, lithofacies Dl has a much higher than random probability of being overlain by lithofacies E because together they have a high positive difference matrix value (0.226).
Table A-4. Difference imatrix.
Lithofacies C Dl D2 D3 EF Row sum
C 0 -0.015 0.032 0.038 0.003 -0.058 0 Dl 0.071 0 -0.194 --0.027 0.226 -0.075 0 D2 0.145 -0.115 0 -0.036 -0.066 0.072 0 D3 0.071 0.011 -0.055 0 -0.075 0.049 0 E -0.343 0.900 -0.186 --0.300 0 -0.071 0 F -0.348 -0.101 0.212 0.296 -0.058 0 0
Results of Table A-4 are presented graphically in Figure 24. APPENDIX 2
Chi-square Test
A chi-square test (Hiscott, 1981) is applied to evaluate the significance of the results derived from the Markov chain analysis.
n
n = rank of the matrix (6 in this study) f.. = the probability matrix from the ith row and the jth column (fCD1 = 0.125 in Table A-2)
s. = sum of the ith row in the transition count matrix 1 (s^ = 24 in Table A-l) r.. = independent trials probability matrix (rCDl = °*140 in Table A-3) ? The calculated^ = 56.68 (1)
Degree of Freedom = the number of non-zero entries in the independent trials probability matrix (Table A-3), minus the rank of the matrix
Degree of Freedom = 30 — 6 = 24 (2)
From the table of chi-square values, the limiting value at 99.9%
confidence level with 24 degrees of freedom = 51.2
56.8 > 51.2
Thus, the results derived from the Markov chain analysis are
highly significant.
97 APPENDIX 3
Described Stratigraphic Sections
The measured sections are presented in numerical order on the following eight pages. Location of the sections are shown in Figure 3.
Unless specified, all sections belong to the Terraba Formation.
LEGEND
SILT & CLAY VERY FINE TO FINE SAM) BIVALVES MEDIUM TO COARSE SAND VERY COARSE SAND E 3 COAL FRAGMENTS -GRAVEL I ULLLl
PELOIDS MASSIVE CROSS-STRATIFIED ✓ LARGE FORAMINIFERA SAMPLE #
NORMALLY GRADED * *\ PLANKTIC FORAMINIFERA
80FT-SEDIMENT DEFORMATION © © DASYCLADACEAN ALGAE
STRATIFIED ^ RHODOLITHS FAULTED
ASH LAYER ALGAE
BURROWED GASTROPODS
LIME MUD
98 99
PY 23
LITHOFACIES LITHOFACIES
70m -
15m - 40m : 3 -20
D3
-14
10m - 60m
4c
-13
2G23 —18
-12 £ E3 -17 -3 5m - 30m J SE3 *Vii
-2 16 '• float
O m 25m - 50m
PY22 -13 LITHOFACIES - 9
2 m - D3
20m
Om PY 41 LITHOFACIES 100 PY 33 LITHOFACIES
PY 35 25m LITHOFACIES
- S3
D3
2 0 m 2 0 m
-14 F -13-2 -13-1 -15— D3
15m 15m
D3
PY 34 LITHOFACIES 1 0 m 1 0 m 4m -5 - 2 -5 -1
M
3-2
-2
Om 5m 5m ) - 7
D2
Om Om 101 PY 44 PY 49 LITHOFACIES LITHOFACIES
*A D3 5m - -4 3m ■ - 2 ___ PY 56 c LITHOFACIES -
D3 Jj-l__ Om
Om
PY 51 1 0 m - LITHOFACIES PY 46 LITHOFACIES C
15m D2
15m 5m
1 0 m
D3 1 0 m Om
a
5m
5m
Om D3
Om-1 PY 57 102 LITHOFACIES PY 66 D2 LITHOFACIES D3 -ti D1
D3 5m
D2
50m
20m Om
PY 67 LITHOFACIES
14m
D2 10m
40m - D2
10m -
J-8 D3
5m -
30m
Om D3 103
PY 68 LITHOFACIES PY 69 PY 70 -2 LITHOFACIES LITHOFACIES D1
D3
2 0 m D1 D1 5m
2 0 m D1
D1
03 15m
D1 Om
1 0 m
1 0 m -
V ✓f
5m
Om
Om 104
PY 71 PY 72 LITHOFACIES a LITHOFACIES
50m - D3
1 0 m .
2 0 m -
Sm - D3
40m I 35?
Om
M m
VVi'fc 1 0 m - M
& PY 74
0 LITHOFACIES 0
M a
5m Claro Member
-3 30m - ■i B *9- M
0m Om 105 PY 75 PY 77
LITHOFACIES LITHOFACIES
35m -l
i p 15m klt
9
20m _ik
30 m - D2 Claro Mentbar < * 0U B 10m - o-"i?
(V^5s&Z, ^ —
\sga
5m 0^<3 ,e.o $ sClaro Mambar A 10m
5 ? D2 i±i^ tvH 4 Brito Fm. Om
■.
D3 Om 106
PY 78 PY 79 LITHOFACIES LITHOFACIES I C
D2 4m J 6m J D2 r 76-4
2m ~C~ D3 D2 3m H Om
Om J APPENDIX 4
Sandston Point Count Data
Point count categories:
DFG (detrital framework grains) = Q+K+P+L+Pyx+A+Op IM (interstitial material) = Carb+Zeol+Phy+Feox+Ortho+Epi Other = Fossil+Epid DFG + IM + Other = 100%
Q = quartz K = potassium feldspar P = plagioclase L (lithic fragments) = Lv+Ls = v+nrfl+Ls Pyx = pyroxene A = amphibole Op = opaque grains F = K + P
v = vitric volcanic grains m = microlitic volcanic grains 1 = lathwork volcanic grains Ls = sedimentary lithics Lv = v + m + 1
Carb = carbonate cement Zeol = zeolite cement Phy = phyllosilicate cement Feox = iron oxide cement Ortho = orthomatrix Epi = epimatrix Fossil = fossil fragment Epid = epidote
107 Table A-5. Modal analysis of sandstone samples of the Terraba Formation.
DFG IM Other Q K P L Pyx A Op v m 1 Ls Carb Zeol Phy Feox Ortho Epl ( 328 93.3 4.3 2.4 0.0 0.0 6.1 86 . 6 0. 6 0 . 0 0 . 0 83 . 5 2.4 0.6 0 . 0 0 . 0 4 .3 0 . 0 0. 0 0. 0 0. 0 318 95.6 3.8 0.6 0.0 0.0 8.8 86.5 0.3 0 . 0 0. 0 8 4 . 6 1.6 0.3 0 . 0 0.3 2.8 0.0 0.0 0 . 0 0 . 6 340 91.5 8.5 0 . 0 0 . 0 0. 0 1 6 . 8 72.4 1.8 0 . 0 0. 6 69.7 2 . 6 0 . 0 0 . 0 0 . 0 0 . 0 8.2 0. 3 O. G 0. 0 303 9 9 . 0 0.3 0 . 6 0 . 0 0. 0 3 6 . 0 61.1 2.0 0 . 0 0. 0 55.8 3.0 2.3 0 . 0 0 . 0 0 . 0 0. 3 0 . 0 0. 0 0 . 0 315 95.2 4.4 0 .3 0. 0 0 . 0 18.1 72.7 3.2 0 . 0 1.3 6 9 . 5 1.9 1.3 0 . 0 0 . 0 0 . 3 0 . 6 0 . 0 0. 0 3.5 307 98.7 0.7 0.7 0.0 0. 0 22.8 75.6 0.3 0. 0 0. 0 68 . 7 4.2 2.6 0. 0 0. 0 0 . 0 0 . 0 0. 0 0.7 0.0 338 8 9 . 6 10.4 0. 0 0. 0 0. 0 1 2 . 1 77.2 0. 0 0 . 0 0. 3 75 . 4 1.5 0. 3 0 . 0 0 . 0 8. 6 0 . 9 0 . 0 0. 0 0.9 361 82.3 15.2 2.5 0 . 0 0 . 0 9.4 71.5 0.0 0.0 1.4 69 . 8 1.4 0.3 0 . 0 15.2 0 . 0 0 . 0 0. 0 0 . 0 0.0 326 95.1 4.6 0.3 0.3 0.0 16.6 77.6 0.6 0 . 0 0. 0 70 . 9 4 . 0 2.8 0 . 0 0 . 0 4. 6 0 . 0 0 . 0 0. 0 0. 0
348 88.8 10.9 0.3 0 . 0 0. 0 18.4 69.3 1.1 0 . 0 0. 0 65 . 5 0. 6 3.2 0. 0 0 . 0 7. 2 3.4 0 . 0 0 . 0 0.3 330 93.3 5.8 0.9 0.0 0 . 0 23.6 67.0 2.7 0. 0 0. 0 60.3 3.9 2.7 0. 0 1.2 4.2 0 . 0 0.3 0. 0 0. 0 331 91 . 0 9.0 0. 0 0 . 0 0. 0 20.5 69.5 0.9 0. 0 0. 0 65 . 6 3.0 0. 9 0 . 0 0 . 0 8 . 2 0 . 9 0 . 0 0. 0 0. 0 357 89.6 10.4 0. 0 0 . 0 0.0 20.0 66.4 3.4 0. 0 0. 0 53.0 6.7 6.4 0 . 0 1.7 7.6 1.1 0. 0 0. 0 0.0 334 90.1 9.6 0. 3 0.3 0. 0 20.4 65.0 3.9 0. 0 0.6 58.1 4.5 2.4 0 . 0 1.8 7.2 0 . 0 0 . 0 0 . 0 0.6 350 86.6 12.6 0.9 0.0 0. 0 16.6 69.1 0.9 0 . 0 0. 0 58 . 0 9.7 1.4 0. 0 12.6 0. 0 0 . 0 0. 0 0.0 0. 0 327 89.6 9.8 0. 6 0 . 0 0. 0 10.4 77.4 1.8 0. 0 0. 0 76.1 0.9 0.3 0.0 0.0 9. 8 0. 0 0 . 0 0 . 0 0. 0 346 87.3 12.7 0. 0 0 . 0 0 . 0 17.3 68 . 2 1.7 0. 0 0 . 0 54 . 9 11.0 2.3 0. 0 5.8 6.6 0 . 0 0 . 0 0.3 0. 0 332 91.9 8.1 0.0 0.0 0 . 0 13.9 76.8 1.2 0.0 0.0 70 . 5 5.7 0.6 0 . 0 4. 5 3.6 0 . 0 0. 0 0. 0 0. 0
342 88.3 11.1 0. 6 0. 0 0 . 0 12.0 75.4 0.9 0. 0 0 . 0 64 . 0 6.7 4.7 0. 0 9.4 1.8 0 . 0 0. 0 0. 0 0.0 354 83.9 16.1 0 . 0 0 . 0 0 . 0 16.7 65.3 2.0 0. 0 0 . 0 56.5 5.1 3.7 0.0 16.1 0 . 0 0 . 0 0 . 0 0. 0 0. 0 389 82 . 0 18.0 0. 0 0.3 0 . 0 9.8 71.2 0.8 0.0 0.0 67.4 2.3 1.5 0. 0 17.7 0. 0 0 . 0 0.3 0 . 0 0. 0 421 73.4 26.6 0. 0 0 . 0 0 . 0 11.4 60.8 1.2 0. 0 0 . 0 57.2 2.1 1.4 0. 0 26.6 0. 0 0 . 0 0 . 0 0. 0 0. 0 392 78.6 21.4 0 . 0 0 . 3 0 . 0 18.1 6C L 2 0. 0 0. 0 0. 0 58.7 1.3 0.3 0.0 21.4 0.0 0 . 0 0 . 0 0. 0 0. 0 376 80.9 19.1 0.0 0.3 0 . 0 17.8 55.3 7.4 0.0 0.0 51.3 2.9 0.8 0.3 0.5 1.3 10.4 2.4 0. 0 4.5 307 98.0 2.0 0. 0 0 . 0 0 . 0 17.3 79.2 1.6 0.0 0.0 67.1 9.1 2.9 0 . 0 0.0 1.3 0.7 0.0 0. 0 0.0 313 97.4 2.2 0. 3 0 . 0 0 . 0 23.6 72.5 1.3 0. 0 0. 0 62 . 9 5.4 4.2 0. 0 0.0 1.6 0. 6 0. 0 0. 0 0. 0 328 92.1 8.0 0. 0 0 . 0 0. 0 16.8 71.0 4.0 0 . 0 0.3 69 . 2 1.8 0.0 0 . 0 7.0 0 . 0 0. 9 0. 0 0 . 0 0 . 0
327 96.0 4.0 0. 0 0 . 0 0. 0 28.1 63.6 4.3 0 . 0 0 . 0 59 . 6 3.4 0.6 0 . 0 0. 0 3.1 4. 3 0 . 0 0. 0 0.3 368 85.9 14.1 0.0 0.0 0 . 0 13.0 58.7 14.1 0 . 0 0. 0 54.3 1.6 2.7 0.0 10.9 1.1 2.2 0. 0 0. 0 0.0 356 85.1 14.9 0. 0 0 . 0 0. 0 6.2 78.9 0.0 0 . 0 0. 0 71 . 9 4.2 2.8 0 . 0 14.0 0 . 0 0. 8 0. 0 0. 0 0 . 0 363 91.7 8.3 0.0 0.0 0 . 0 28.7 52.1 8.5 0.3 2.2 47 . 4 3.3 1.4 0 . 0 0.0 0.0 8.3 0. 0 0. 0 0.0 426 75.4 24.6 0. 0 0 . 0 0. 0 49.5 25.4 0.0 0 . 0 0.5 24 . 9 0. 0 0.5 0.0 20.4 0.0 4.0 0.2 0. 0 0.0 347 88 . 2 11.5 0.3 0.6 0.0 32.0 54.5 0.9 0. 0 0.3 43 . 5 6.6 4.3 0. 0 1.4 0 . 6 5.8 0.3 0. 0 3.5 318 97.2 2.8 0. 0 0 . 6 0. 0 16.7 79.6 0.3 0 . 0 0. 0 68 . 9 8.2 2.5 0 . 0 0.0 0.9 1.6 0.0 0.3 0.0 345 88.1 11.9 0. 0 0 . 0 0 . 0 35.7 51.3 0.6 0 . 0 0. 6 49 . 6 0. 6 1.2 0.0 0.9 0.6 0.0 3.2 2.0 5.2 332 94.9 4.8 0.3 0 . 0 0. 0 2 5 . 0 69.3 0.3 0 . 0 0.3 59 . 9 7.2 2.1 0. 0 2.1 1.8 0.6 0. 0 0 . 0 0.3
379 79.9 20.3 0. 0 12.9 0. 0 43 . 8 23.0 0.0 0 . 0 0. 0 20.1 1.3 1.6 0.0 0.0 18.5 1.8 0.0 0.0 0.0 359 93.0 4.2 2.8 0.0 0. 0 30.1 60.4 1.7 0 . 0 0. 8 59.1 0. 6 0. 8 0. 0 0.8 0 . 0 2.8 0.0 0.0 0.6 379 81 . 0 18.2 0. 8 0 . 0 0. 0 3 4 . 0 36.9 9.0 0.3 0. 8 25.1 8.2 3.7 0.0 5.5 0. 0 12.7 0.0 0.0 0.0 548 58.8 10.9 30.3 0 . 0 0 . 0 10.4 4 8 . 0 0. 0 0 . 0 0. 4 37 . 6 4.7 0.7 4.9 6.0 0. 2 0 . 0 0. 0 4.7 0.0 328 97.3 1.8 0.9 0.0 0. 0 23.5 73.2 0.0 0 . 0 0.6 68.3 4.6 0.3 0.0 0. 0 0.3 1.5 0. 0 0.0 0.0 322 96.0 3.7 0.3 0 . 0 0 . 0 21.1 73.3 1.6 0.0 0.0 6 5 . 2 4.7 3.4 0.0 0.0 3.7 0.0 0.0 0.0 0.0 329 93.3 6.7 0. 0 0 . 0 0. 0 24.3 67.2 1.8 0.0 0.0 50.8 12.8 3.6 0.0 0.0 6.4 0.3 0. 0 0. 0 0.0 320 97.5 2.5 0. 0 0. 0 0 . 0 20 . 0 76.9 0.6 0.0 0. 0 66 . 9 6.3 3.8 0 . 0 0. 0 2.5 0 . 0 0 . 0 0 . 0 0.0 109
Table A-6. Recalculated values from Table A-5. Lv - volcanic llth ic s PyxOp « pyroxene + opaque grains. Q+F+L “ 100%, Lv+PyxOp+P ” 100%, v+nrfl » 100%, v' + P+PyxOp + m+1 « 100%.
Sample Q FL Lv PyxOp P V m 1 v' p+pyxOp nri-1 number
PY-7 0.0 6.6 93.4 92.8 0.7 6.5 96.5 2.8 0.7 89.5 7.2 3.3 PY-8 0.0 9.2 90.8 90.5 0.3 9.2 97.8 1.8 0.4 88.5 9.5 2.0 16-1 0.0 18.8 81.2 79.1 2.6 18.3 96.3 3.7 0.0 76.2 20.9 2.9 18-1 0.0 37.1 62.9 61.7 2.0 36.3 91.4 4.9 3.8 56.3 38.3 5.3 19-4 0.0 19.9 80.1 76.3 4.7 19.0 95.6 2.6 1.7 73.0 23.7 3.3 21-2 0.0 23.2 76.8 76.6 0.3 23.1 90.9 5.6 3.4 69.6 23.4 6.9 21-4 0.0 13.6 86.4 86.1 0.3 13.5 97.7 1.9 0.3 84.2 13.9 1.9 23-11 0.0 11.6 88.4 86.9 1.7 11.4 97.7 1.9 0.4 84.8 13.1 2.0 33-3-1 0.3 17.5 82.1 81.9 0.6 14.5 91.3 5.1 3.6 74.8 18.1 7.1
33-3-2 0.0 21.0 79.0 78.0 1.3 20.7 94.6 0.8 4.6 73.8 22.0 4.2 33-5-2 0.0 26.1 73.9 71.8 2.9 25.3 90.0 5.9 4.1 64.6 28.2 7.1 33-6 0.0 22.8 77.2 76.4 1.1 22.6 94.3 4.3 1.3 72.1 23.6 4.3 33-7 0.0 23.1 76.9 74.1 3.8 22.2 80.2 10.1 9.7 59.4 25.9 14.7 33-9 0.3 23.8 75.9 72.3 5.0 22.7 89.4 6.9 3.7 64.7 27.7 7.7 33-10-2 0.0 19.3 80.7 79.9 1.0 19.1 83.4 14.0 2.1 67.0 <• 20.1 12.9 33-13 0.0 11.8 88.2 86.3 2.0 11.6 98.4 1.2 0.4 85.0 13.7 1.4 33-14 0.0 20.3 79.7 78.1 2.0 19.9 80.5 16.1 3.4 62.9 21.9 15.2 33-17 0.0 15.3 84.7 83.6 1.3 15.1 91.8 7.5 0.8 76.7 16.4 6.9
33-20 0.0 13.7 86.3 85.4 1.0 13.6 84.9 8.9 6.2 72.5 14.6 12.9 33-21 0.0 20.3 79.7 77.8 2.4 19.9 86.6 7.8 5.6 67.3 22.2 10.4 34-3-1 0.3 12.0 87.7 87.1 0.9 11.9 94.6 3.2 2.2 82.4 12.9 4.7 34-3-2 0.0 15.8 84.2 82.8 1.6 15.5 94.1 3.5 2.3 78.0 17.2 4.9 42-2 0.3 23.1 76.6 76.9 0.0 23.1 97.5 2.1 0.3 74.9 23.1 2.0 44-4 0.3 24.3 75.4 68.5 9.3 22.2 93.2 5.3 1.4 63.9 31.5 4.6 46-2 0.0 17.9 82.1 80.7 1.7 17.6 84.8 11.5 3.7 68.4 19.3 12.3 47-3 0.0 24.6 75.4 74.4 1.3 24.3 86.8 7.5 5.7 64.4 25.6 9.8 48-2 0.0 19.1 80.9 77.2 4.6 18.2 97.4 2.6 0.0 75.2 22.8 2.0
48-18 0.0 30.7 69.3 66.2 4.5 29.3 93.8 5.3 1.0 62.1 33.8 4.1 53-1 0.0 10.6 89.4 68.4 16.5 15.2 92.6 2.8 4.6 63.6 31.6 5.1 57-10 0.0 7.3 92.7 92.7 0.0 7.3 91.1 5.3 3.6 84.5 7.3 8.3 63-3 0.0 35.5 64.5 56.9 11.7 31.3 91.0 6.3 2.6 51.8 43.1 5.1 63-6 0.0 66.1 33.9 33.6 0.6 65.7 98.1 0.0 1.9 33.0 66.4 0.6 66-2 0.7 36.8 62.9 62.2 1.3 36.5 79.9 12.2 7.9 49.7 37.8 12.5 67-1 0.6 17.2 82.1 82.4 0.3 17.3 86.6 10.3 3.2 71.3 17.6 11.1 68-2 0.0 41.0 59.0 58.2 1.3 40.5 96.6 1.1 2.3 56.3 41.8 1.9 69-2 0.0 26.5 73.5 73.0 0.6 26.3 86.5 10.4 3.0 63.2 27.0 9.8
70-1 16.2 55.0 28.8 34.4 0.0 65.6 87.4 5.8 6.9 30.0 65.6 4.3 71-1 0.0 33.2 66.8 65.0 2.7 32.3 97.7 0.9 1.4 63.5 35.0 1.5 74-5 0.0 46.0 52.0 45.8 12.1 42.2 67.9 22.1 10.0 31.0 54.2 14.7 78-2 0.0 17.8 82.2 80.0 0.7 19.3 87.3 11.0 1.7 69.8 20.0 10.2 NN 0.0 24.3 75.7 75.2 0.6 24.1 93.3 6.3 0.4 70.2 24.8 5.0 NA1-2 0.0 22.4 77.6 76.4 1.6 22.0 89.0 6.4 4.7 68.0 23.6 8.4 NA4-3 0.0 26.6 73.4 72.0 2.0 26.1 75.6 19.0 5.4 54.4 28.0 17.6 NA4-4 0.0 20.6 79.4 78.8 0.6 20.5 87.0 8.1 4.9 68.6 21.2 10.3 VITA
Peter Bee-Deh Yuan, the son of Chi-Fu Yuan and Wen-Chlng Soong
Yuan, was born on October 6, 1948, in Hankow, China. He was raised in
Pingtung, Taiwan where he attended elementary and high schools. He received a B.S. degree in Earth Sciences from National Cheng Kung
University in 1973. After two years of military service in the army of the Republic of China, he began graduate studies at Carleton University,
Ottawa, Canada, from which he received a M.S. degree in geology in 1979.
In the same year, he began a Ph.D. program at Louisiana State Universi ty.
In 1975, Peter married Kim-Jean Chow, daughter of Mr. and Mrs. Ray-
Liang Chow of Taipei, Taiwan. They have two daughters, Jia-Luh and Jia-
Chyi, born on May 19, 1979, in Baton Rouge, Louisiana. The Yuans now reside in Tainan, Taiwan, where Peter is employed by National Cheng Kung
University as an assistant professor.
1 1 0 EXAMINATION AND THESIS REPORT
Candidate: Peter B. Yuan
Major Field: Geology
Title of Thesis: Stratigraphy, Sedimentology, and Geologic Evolution of Eastern Terraba Trough, Southwestern Costa Rica
Approved:
Major Professor a: lairm
Dean of the Graduate School
EXAMINING COMMITTEE:
Date of Examination:
May 9, 1984