Pelagic redbeds in the Devonian of Germany — deposition and diagenesis

Sedimentary Geology, 25 (1980) 231--256 © Elsevier Scientific Publishing Company, Amsterdam -- Printed in The Netherlands

PELAGIC REDBEDS IN THE DEVONIAN AND DIAGENESIS *

OF GERMANY

231

-- DEPOSITION

WOLFGANG F R A N K E and J O S E F PAUL

Geologisch-Pala'ontologisches Institut und Museum der Universit~t GSttingen, D-34 G6ttingen (F.R.G.) (Received May 7, 1979; revised and accepted September 6, 1979)

ABSTRACT Franke, W. and Paul, J., 1980. Pelagic redbeds in the Devonian of Germany -- deposition and diagenesis. Sediment. Geol., 25: 231--255. Red pelagic mudstones are a conspicuous facies type of the Variscan basins in Europe, where they occur chiefly in the Famennian. They are restricted to basinal sites without major clastic influx: in areas marginal to turbidite fans, in starved intra-geosynclinal basins, or on the slopes of rises. Thin silt layers, intercalated among the red shales, are interpreted as distal or lateral turbidites. Pelagic body fossils are extremely rare. Endobentonic trace fossils, indicative o f outer fan and trough environments, are proof of aerated bottom water. Ferric iron, from which hematite is formed during diagenesis, is provided by continental weathering, and is mainly bound to the clay fraction. Red colouration in mudstones requires a minimum of about 1.5% hematite, as calculated from chemical analyses. The content of pigment may be taken below the critical level by dilution with non-red silicate or carbonate. Conversely, enrichment of the clay fraction, at places far from the source area for example, enhanced a red colouration. Drab colouration is mainly due to reduction of iron, caused by oxidation or organic matter during early diagenesis. Reduction continued during later diagenesis and tectonic deformation, as demonstrated by colour-boundaries which cut through the bedding and partly follow structured patterns. Where the sediment was permeable (e.g., prior to compaction or along structural surfaces), the reduced iron was partly removed. The red pigment could only survive where organic supply to the sediment was too low to consume the available oxygen. The deficit of organic matter in the Famennian sediments is attributed to oligotrophic conditions. These could have been brought about by decrease in continental afflux from the Caledonian source areas, which, by the late Devonian time, were much reduced by erosion and also progressively more limited in areal extent as marine transgression advanced. Oligotrophy is seen as one of the main factors governing the origin of all kinds of redbeds: terrestrial, fluviatile, estuarine, lagoonal-hyperhaline, nearshore-marine, as well as (in the Famennian) pelagic-bathyal.

* This work forms part of the research program of the Sonderforschungsbereich 48 'Entwicklung, Bestand und Eigenschaften der Erdkruste, insbesondere der Geosynklinalr~/ume', University of GSttingen (Federal Republic of Germany).

232 INTRODUCTION The vast majority of ancient redbeds have been deposited in environments on or close to land. This is also valid for the sediments of the Variscan Geosyncline in Europe: the deposits of the Old Red Sandstone Facies fill up a considerable portion of various Devonian troughs. Next in rank, in terms of thickness are the red clay- and siltstones of early to middle Devonian age which were deposited in inner shelf areas transitional between the Old Red Sandstone environment on the north and the open sea to the south and southeast. The present paper deals with pelagic redbeds. Their volume is negligible compared with other facies. However, they are widespread in the Upper Devonian in southwest England and in Germany (Rhenohercynian and Saxothuringian Zones of the Variscan Belt in Europe). Besides, they form an isochronous marker bed in the Famennian of the Rheinisches Schiefergebirge, and thus seem to reflect some special hydrographic or palaeogeographic situation in the geosynclinal basin. This problem is here approached with the aid of evidence on the regional distribution of pelagic redbeds in space and time, and also from detailed geochemical and sedimentological analyses of samples from selected sites. The study of the Famennian case leads to some conclusions about deposition and diagenesis of marine redbeds in general. REGIONAL AND STRATIGRAPHIC DISTRIBUTION OF PELAGIC REDBEDS

Rheinisches Schiefergebirge Fig. 1 gives a map and a generalized picture of the Devonian topography as it existed in the region east of the River Rhine. The topography is shown as a section which runs from the northwestern shelf {adjacent to the Old Red Sandstone facies) to the metamorphic uplift on the opposite of the geosynclinal basin ('Mitteldeutsche Schwelle', Brinkmann, 1948). The figure illustrates an early Famennian situation, but the same picture remains generally valid through to the Tournaisian, where the pelagic sediments to be discussed here were succeeded by a uniform sequence of Tournaisian/Vis~an black shales and cherts. Information on the general depositional history has been published in recent years by Meischner {1971), Krebs {1971), and Franke et al. {1978). A survey of the Famennian clastics has been published by Einsele {1963).

Northwestern shelf (neritic). The sequence is composed of shallow marine sandstones and shales which first appear at about the base of the Famennian and continue into the basinal Carboniferous. Neritic red shales are found only in the Boulonnais of northern France (Delattre et al., 1973).

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sandstones from the NW

greywackes from the SE

neritic I hemi - pelagic I pelagic Fig. 1. Palaeogeographic map and simplified morphological cross-section of the Scbiefergebirge east of the River Rhine during the lower Nehden-Stufe (basal Famennian). Outcrop of Palaeozoic rocks is within the continuous line. Black in the map: outcrop of Upper Devonian; black in profile: sites of Famennian redbed deposition. Brick ornament (map and profile): Givetian/Frasnian reef complexes. Stippled rim on map: outer margin of basinal Nehden Sandstone; stippled in profile: Nehden Sandstone. Coarser dots: greywackes. Arrows: main current direction of Nehden Sandstone (from Einsele, 1963). Sampling localities: E f (Effenberg), Ka (Kahlenberg), L i (Liesen), B r (Braunau), N a u (Nauborn).

234

Drowned Middle Devonian shelf (hemi-pelagic). During the Givetian, an extensive carbonate platform was built up over a foundation of neritic clastics. Isolated Givetian to Frasnian reef complexes growing upon the platform created highs and inter-reef lows (Meischner, 1964; Krebs, 1974). Since the Famennian and Tournaisian (this applies to the whole of the Schiefergebirge) were periods of geotectonic inactivity (Sonderforschungsbereich 48, 1974; Franke et al., 1978), these topographic features persisted into the early Carboniferous. During the Famennian, condensed carbonate sequences were laid down on the old reef tops. Time-equivalent clastics, derived from the northwestern shelf, were deposited in inter-reef basins. The sandstones are grey or greenish. Red colour is generally restricted to shale units between the major sandstone bodies and is most c o m m o n in the basal part of the Hemberg stage (middle Famennian; see Ziegler, 1971, for ammonoid and c o n o d o n t zonation). The 'Hemberg-Rotschiefer' may attain a thickness of more than 100 m, whereas the other red shale horizons range between 1 and 10 m. The whole clastic sequence of the Famennian reaches a maximum of 600--700 m.

Basin (pelagic). The topography of the basin south of the drowned shelf is differentiated only by a few Givetian occurrences of volcanics and some minor basement rises near the former shelf margin (not shown in Fig. 1). The thicknesses and the stratigraphic succession of red shales and sandstones are essentially the same as in the hemi-pelagic area. Red shales also occur on some rises, which were below the level of condensed carbonate sedimentation, but nevertheless above the paths of sand transport.

Southeastern basin and rise system (pelagic). The southeastern Schiefergebirge (Dill Syncline, H5rre-Kellerwald-Zone, Lahn Syncline) is characterized by an intricate pattern of narrow rises and intervening basins, which are arranged parallel with the later-developed fold axes (SW--NE). Some of the rises are basement rises, elsewhere they are formed by Givetian volcanics. The relief is accentuated by the growth of reef complexes on volcanic highs. The northwestern succession of red shale and sandstone bodies can be traced southward as far as the Dill Syncline, where the sands were finally barred by the palaeo-relief. At the same time, the 'Mitteldeutsche Schwelle' in the south (which was later to become the source area of the Carboniferous flysch) was already experiencing uplift and was shedding greywacke material into some of the southern basins (Meischner, 1968; Bender, 1978). Some red shale is intercalated among these greywackes. The majority of the narrow southern basins are devoid of any coarsegrained terrigenous sedimentary rocks. It is in these starved basins that the deposition of red shales became d o m i n a n t (Rietschel, 1966; Goldmann, 1967; Schneider, 1969). In some places there are first occurrences in the Givet stage. Red shales are often found in the Adorf Stage and they dominate the basinal facies t h r o u g h o u t the Famennian, up to the Devonian/Carboniferous

235 boundary. There are even some reports of upper Tournaisian and basal Vis~an red shales. The area between the Lahn Syncline and the 'Mitteldeutsche Schwelle' in the south is occupied by a broad belt of outcrop of older rocks, so that there is no information on the southern margin of the Famennian basin.

Harz Mountains The Palaeozoic in the Harz Mountains forms a northeastern continuation of what is seen in the Schiefergebirge and corresponds in its essentials, with the basin and rise system described earlier (H. Schmidt, 1931). In consequence, the distribution of pelagic redbeds is nearly identical to that of the Schiefergebirge. The only notable exception is the absence of north-derived Famennian sands, none of which reached that far to the east. Details may be found in papers by Reichstein (1955), Stopped and Zscheked (1971), Meischner and Schneider (1971), and in the review published by Mohr (1978).

Southwest England As previous authors have observed (Goldring, 1962; Matthews, 1977), the Devonian of Southwest England has very much in c o m m o n with the Harz Mountains and the Rheinisches Schiefergebirge, and may be regarded as a western part of the Rhenohercynian Zone. Pelagic red shales are c o m m o n in Cornwall. Their deposition probably started during the Frasnian, and, in the P l y m o u t h area, continued t h r o u g h o u t the Famennian (Gooday, 1974). Red slates on the west coast have an early to middle Famennian age (Beese, 1977). The wide stratigraphic range of the pelagic redbeds, as well as their general palaeogeographic setting, are reminiscent of the southern Schiefergebirge.

Saxothuringian Zone The Saxothuringian Zone (Kossmat, 1927), south of the 'Mitteldeutsche Schwelle', contains a suite of geosynclinal sediments and volcanics similar to those of the Rheinisches Schiefergebirge and Harz (Rhenohercynian Zone of Kossmat). Red shales occur locally in the Adorf Stage and are most frequent in the upper Nehden and Hemberg Stages (Pfeiffer, 1968). In some places, there is an admixture of volcanic ashes (Trusheim, 1964). Some aspects of colouration are discussed in Schlegelmilch (1968).

Discussion The distribution of pelagic red shale in space and time leads to the following generalization: (1) The red colouration is restricted to shales and silty shales. It disappears

236

in the vicinity of sandstone bodies. Red shales achieve their longest stratigraphic range in basins from which the coarser clastics were barred. The red shales again disappear up the slope of intra-basinal rises, where they are replaced b y pelagic limestones which are, in the Schiefergebirge, generally non-red. In b o t h the sandstone and the limestone environments, the red pigment was either diluted by non-red minerals to ineffective levels of concentration, or was destroyed by chemical reduction. Both possibilities will be discussed in closer detail below. (2) Although red shales were deposited at various times from the Givetian to the beginning of the Vis6an, most reports are from the early and middle Famennian. Even in areas where the distribution of red shales is not subject to the influence of intercalated sandstone bodies (northwestern Harz Mountains), redbeds are restricted to the uppermost Nehden and lower Hemberg Stages. Possible reasons for this will be discussed later in the paper. LITHOLOGY

The prevailing sediment of the red shale successions is a mixture of silt and clay. Silt-size quartz and mica may be found dispersed through the sediment, b u t are for the most part arranged in parallel laminae, at separations of the order of 0.1--1 mm. The amount of silt is generally low on rises and in restricted basins. There is a small and variable carbonate c o m p o n e n t in the matrix. It becomes dominant on rises, where it is not diluted by clastic sedimentation. In slope and rise successions with a high carbonate content, the carbonate is concentrated in nodules and layered bodies whose origin is diagenetic (see Fig. 2). Models of diagenetic precipitation have been proposed b y Griindel and Rbsler (1963, Devonian) and Jenkyns (1974, Mesozoic).

TABLE I N o n - c a r b o n a t e minerals in t h e r e d shales f r o m Liesen, same samples as in Fig. 2 ( i n f r a r e d s p e c t r o s c o p y , all d a t a in %) Li

1

2

3

4

5

6

7

Rhythm member Colour

A

B

C

C

A

B

B

green

purple

red

green

green

green

red

Quartz Feldspar Chlorite Illite

42 8 16 29

26 6 16 50

32 7 11 45

31 5 16 46

35 8 15 38

29 6 18 41

29 6 15 48

--

1

2

--

--

--

1

Hematite

237

EFFEI .....

LIESE

4.!

1_Fig. 2. T w o typical slabs o f pelagic red shales w i t h i n t e r c a l a t i o n o f diagenetic n o d u l a r carb o n a t e (in black, E f f e n b e r g ) a n d silt-layers (Liesen). N u m b e r s : a n a l y z e d samples. A ; silt, B: clay, C: silty clay. D r a b c o l o u r s s h o w n in w h i t e {except for c a r b o n a t e ) , r e d c o l o u r s s h a d e d grey. Scale b a r r e p r e s e n t s 10 cm.

Towards the basinal sandstone bodies, silt layers become thicker and closer spaced. Many silt-bands are immediately overlain by a thin layer of pure clay. The silt/clay couplets are intercalated among the normal, silty claystone to produce repetitions of a rhythmic sequence (see Fig. 2) :

9

10

11

12

14

15

16

17

18

19

C

A

B

C

B

C

A

B

C

A

red

green

red

red

red

red

green

red

red

green

30 8 11 46 2

32 7 14 39 --

30 5 13 49 2

30 6 12 45 3

24 4 16 52 2

29 6 10 49 3

34 8 16 36 1

27 2 13 54 3

29 6 11 48 3

34 8 17 35 --

238 C : silty claystone, laminated, B: pure claystone, structureless, and A: siltstone, parallel lamination or lenticular cross lamination, carbonate- cemented. The petrographic composition of some typical rhythms (figured in Fig. 2) is given in Table I. There is a positive correlation between the thicknesses of the A and the B members, and where the A m e m b e r (silt) is not distinguishable in the field, the B m e m b e r is significantly thinner than in rhythms that are complete. In such cases, an A m e m b e r might be represented by the t o p m o s t silt-lamina of the underlying C member. As thin sections and X-ray diffraction diagrams show, the staining mineral is hematite, which is present as irregular clusters or coatings throughout the rock.

Discussion All of the evidence suggests that the A (silt) and B {pure clay) members represent syngenetic couplets introduced into a basin whose background sedimentation was silty (C). We tentatively interpret the A/B couplets as distal turbidites, deposited as peripheral features of the main Famennian sandstone bodies. The more or less abrupt transition from A to B would, in a turbidite model, imply a break in the grain-size spectrum of the suspension. Probably, part at least of the clay forming the B member is not derived from the shelf, b u t has been eroded from the basin floor during the flow of the turbidity current. A definite conclusion on this point would emerge only from the sedimentological analysis of the coarser clastics, which is not the purpose of this paper. The suggestion that turbidites are present among the Famennian sandstones is still controversial (Einsele, 1963); b u t whatever their origin, the ABC rhythms must be taken into account, because they are characteristic of the basinal facies and have strongly influenced the colouration of the red shales (see page 240). FAUNA The record of b o d y fossils in the red shales is very poor. Goniatites are extremely rare. Conodonts are few in number, b u t are regularly found. The most frequent fossils are entomozoid ostracodes (hence the term 'Cypridinen-Schiefer' in older German literature), which are sometimes found crowded on the bedding planes. They are regarded as nektonic (Rabien, 1954). The diversity of the ostracod fauna is low. Some thin, small bivalves (Posidonia venusta) may have been attached to floating objects (H. Schmidt, 1931, 87). The only benthonic b o d y fossils are exuvia of phacopid trilobites, often found in the Salter m o d e of preservation, which indicates quiet water (Richter, 1937).

239

At many localities trace fossils are the only evidence that benthonic life was possible. We have found the following types (see Fig. 3): (1) Silt-filled tubes, often paired, at various angles with bedding, a b o u t 2 mm in diameter, and with a distance of a b o u t 1 cm between the paired tube endings (Spreiten-burrows of uncertain affinity). (2) Radiating burrows of the Chondrites type. (3) Irregular meanders (Helminthopsis). (4) Regular meanders (Helminthoida type) are the most c o m m o n trace fossils. Most are small {length of loops between 2 and 5 cm). A large t y p e (rarely found) has loops a b o u t 2 mm wide and 10 cm long. (5) A regula.r meander with loops bent into right angles, and short ridges on either side of a central groove (? Nereites). (6) A large and very irregular form of Palaeodictyon, with polygons sometimes of dimensions exceeding 10 cm in diameter and tubes up to 1 cm wide. These trace fossils, except for the paired tubes and Chondrites, have been found in almost every case at the base of the B-member (pure clay), on top of the silt-bands ('A'), or on t o p of the laminated silt/clay-member ('C'), where a miniature silt-band might be assumed (see Fig. 2).

2

3 -

Fig. 3. Trace-fossils from Famennian red shales. 1: Chondrites, 2: Helminthopsis, 3--6: Helminthoida, 7: ? Nereites, 8, 9: Palaeodictyon (polygons elongated due to tectonic deformation). Scale bars represent 10 cm.

240

Discussion As several previous authors have pointed out, the b o d y fauna of the red shales is clearly pelagic. It is n o t e w o r t h y that the same characteristic combination of pelagic redbeds and entomozoid ostracodes reappears in the Upper Carboniferous of the Cantabrian Mountains of Spain (Becker et al., 1975; Becker, 1976). The preferred occurrence of trace fossils at interfaces between layers of contrasting grain size might be an effect of selective preservation, or due to preferential splitting of the shale. On the other hand, the trace-producing animals might have seletively fed upon the coarse-grained layers, where there might have been organic matter introduced into the basin along with the terrigenous debris (see below). On the whole, the ichnofauna closely resembles the 'Nereites' facies of Seilacher (1967), which characterizes deep-water environments. This is consistent with the sedimentological data and the record of b o d y fossils (see above). The ichnofauna of the pelagic red shales is significantly different from that of the 'Nehden-Sandstein', deposited in the northwesternmost part of the same upper Devonian basin (Bandel, 1973), which contains Cruziana and lacks Helminthoida and Palaeodictyon. Since Bandel's fauna has been found near the source area of the basinal sandstones, and the siltstone bands within the red shales represent distal or lateral margins of sandstone bodies, the difference in faunal composition may be explained by the model of Crimes (1977), who found different trace-fossil assemblages in different parts of the same submarine fan. Our 'red shale assemblage' would, in Crimes' table 4, range among the 'outer fan' and 'trough' associations (Palaeodictyon, Helminthopsis, Helminthoida). This fits with our sedimentological interpretation of the siltstone layers. As the occurrence of trilobites and trace fossils would indicate, the bottom water of the Famennian basin was aerated. COLOURATION

The shales appear as an irregular alternation of red and grey or green units, each measuring some dm or m. In the red portions, there are reduction spots up to 1 cm in diameter, which may contain small calcareous nodules, indeterminable fragments of fossils, or c o n o d o n t s as nuclei. Thin silt layers (the A m e m b e r of the sedimentary rhythm) can appear as green bands, from which the green colour extends upward and downward b e y o n d the limit of the silt-size grains, so that the colour boundary lies in the 'B' above and the 'C' below (see Fig. 2). The remaining, upper part of the B-member is reddish-brown (2,5 YR 3/4 of the Munsell soil-color chart), and, in fresh exposures is clearly distinct from the reddish-grey (10 R 3/1) of the C member.

241

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0

:

10

20

..............

30 m

. . . . . . . . . . . . . . . . . . .

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Fig. 4. Reduetion o£ red pigment around thrust faults. Drab colours in white, red colours shaded grey. Section along a football field at Nauborn, S of Wetzlar.

Green bands occur on rises also, where they are found in some of the carbonate-rich layers (see Fig. 2). Carbonate nodules sometimes enclose the colour boundary. Since a carbonate concretion, once grown, is n o t likely to be penetrated by pore-fluids which could effect a change of colour, it can be assumed that the diagenetic development of carbonate either entirely postdated or else continued later than, the development of the differential colouration. The b o u n d a r y between the major green and red units often cuts across the stratification. The colour b o u n d a r y may be oblique to bedding, or m a y follow structural features such as joints, cleavage, or thrust faults (see Fig. 4). Along all colour boundaries there is a transitional mm of greyish purple. Local reddening near the Permian land surface, and bleaching below humic soils are reminders that weathering, too, has an influence; but that question will not be further discussed in this paper. Discussion

Although it was possible for colour boundaries to change position during or after tectonic deformation, the most obvious colour limits are stratabound. Small-scale observations are in complete accord with palaeogeographic findings in that the red colour disappears in the neighbourhood of coarser clastic or calcareous bodies of whatever size. It is an observable fact that the green spots and bands occur in a red matrix (and n o t vice-versa). From this one infers that red is the original colour and it was locally obliterated by reduction of the iron pigmentation. GEOCHEMICAL DATA

The results of the geochemical analyses are compiled in Table II. Since the red pigment is hematite, special attention has been given to analysis of iron and its states o f oxidation. It is evident from the plot of ferric against ferrous iron {Fig. 5), t h a t neither the total iron c o n t e n t nor the ratio ferric/ferrous iron is critical for the colouration of the rock. The only diagnostic value is the minimum c o n t e n t of ferric iron required for redness.

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Fig. 5. Ferric plotted against ferrous iron. All red samples have more than 2% ferric iron. Black symbols: red, half-black: purple, open symbols: drab samples.

In the several examples treated, this value ranges between 2 and 3%. A similar minimum, around 2% ferric iron, can be retrieved from data published b y several authors, who have described redbed-environments of various kinds and of different ages (see Fig. 8). It should, however, be stressed again here, that these figures refer to claystones and silty claystones only (and their altered equivalents in orogenic belts). In coarser clastics, the red pigment often occurs as a thin grain-coating, and much less iron is needed to produce a red colour (Hartmann, 1963). Iron was also determined from sodium-dithionite extracts -- a method used in soil science to estimate the largest portion of the iron fraction n o t incorporated into silicate lattices. There is a positive, highly significant correlation between whole-rock ferric iron and dithionite-extracted iron (Fig. 6). This supports the idea that the ferric iron content is a measure of the a m o u n t of (leachable) hematite. The regression line of Fig. 6 indicates an excess of ferric over extracted iron. This might suggest that, under the analytical conditions obtaining, hematite is not leachable quantitatively. Only part of the ferric iron determined in whole-rock analyses is present as hematite. At 0% of extracted iron (Fig. 6), there is still a b o u t 0.4% of ferric iron present. This remaining half percent is probably fixed into clay minerals. A similar level of ferric iron in the lattice of clay minerals may be deduced from Figs. 5, 6 and 7. The green samples (which have suffered from natural leaching processes) never have less than 1% of ferric iron. This figure has to be substrated from the 2% minimum in Fig. 8, to obtain the minimum of 'free' ferric iron required for red colouration: 1% Fe 3÷ (or approx. 1.5% hematite).

246 Feextr

3-

°/o

2

1

0

1

2

3

4

5

6

7

°/o Fe 3.

Fig. 6. Positive correlation of ferric iron and sodium-dithionite-extracted iron. Symbols as in Fig. 5.

In order to obtain the a m o u n t of ferric iron, which is imported with the terrigenous fraction, the m e a n ratio ferric iron/non-carbonate fraction was calculated for the red samples of the different sections. The highest values were obtained for Nauborn and Braunau (4.1% and 3.6% ferric iron in the

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Fig. 7. Ferric and ferrous iron in red/drab sampling couplets, taken from thin shale bands across which the colour boundary runs. The purple samples have been taken from the colour boundary itself. The tie lines indicate changes in the balance of ferric and ferrous iron brought about by reduction (and part removal) of iron at different stages of diagenesis. Colour-boundaries: a) reduction spots, b) reduction-bands, c) oblique to bedding, no specification possible, d) parallel to cleavages, e) along thrust-faults, f) along joints. Other symbols as in Fig. 5.

247 0 U. Cretaceous

Mexico

U. Cretaceous

AIps

L. Triassic U. Devonian

1

3

4 °/0 Fe 3• delta

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L.-M. Devonian, div.

2

A

Mc Bride 1974

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Faupl & Sauer 1978

supratidal plain

Paul & Franke 1977

pelagicbathyal

this paper

I IIIII

I

I

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I

_ _ ] --1

I

I

Reichstein 1955 /

-

-

~ I

nearshore~estuarine or

Franke & Paul unpubl.

mixed - marine ?

Schulz Dobrick 1975

Fig. 8. F e r r i c i r o n c o n t e n t o f various r e d m u d s t o n e e n v i r o n m e n t s . R i g h t - h a n d e n d o f w h i t e c o l u m n : m a x i m u m ferric i r o n - c o n t e n t o f d r a b samples, l e f t - h a n d e n d o f b l a c k c o l u m n : m i n i m u m ferric i r o n c o n t e n t o f red samples. All red samples e x c e e d a critical l i m i t o f a b o u t 2% o f ferric iron.

non-carbonate portion). Both sections are situated in these basins which were barred from the coarser clastic influx. To establish quantitative control of the reduction processes, several shale bands of homogeneous lithological composition have been traced and analysed across the green/red colour boundary. The iron budgets of several green/red couplets, and the nature of the respective colour boundaries are shown in Fig. 7. There are different cases of reduction, which are discernible from the behaviour of the reduced (ferrous) iron. In reduction spots and reduction bands (around silt layers), the reduced iron is more or less quantitatively removed from the system (Braunau, Liesen). Where the colour boundary runs at an angle to bedding, w i t h o u t following any identifiable sedimentary or structural features, a great portion of the reduced iron is retained within the system, and is probably fixed in the carbonate minerals (Kahlenberg, Liesen). In the vicinity of thrust faults (Nauborn, see Fig. 4), only a minor part of the reduced iron is held back. Near vertical joints (Effenberg), all reduced iron is removed, and even the stock of primary ferrous iron is attacked -- this might be due to the dissolution of ferruginous carbonate. The cases cited above correspond to different stages of diagenesis and very low grade metamorphism. The budgets are based on the assumption that no iron has been added to the system after deposition. The distribution of m a n g a n e s e is totally independent of the colour of the sample. There is a clear positive correlation with carbonate content, and an equally positive correlation between the manganese and calcium extracted with sodium dithionite. Evidently, the bulk of manganese is fixed into carbonate minerals.

248

Organic carbon (mean value of 0.04% in the Liesen section) is low compared with the average for Rhenohercynian shales (0.24%: Schulz-Dobrick, 1975, table 5). The c o n t e n t of the green samples seem to be slightly higher than that of the red ones (0.05 and 0.04% respectively). The c o n t e n t of phosphate (mean of 1500 ppm) corresponds with the average of values in Schulz-Dobrick's table 5. MODEL OF DEPOSITION AND DIAGENESIS

Iron is transported and (in most cases) deposited under oxidizing conditions. It occurs as colloidal and particulate oxides and hydroxides, or adsorbed to clay minerals. Hematite is formed by diagenetic alteration of hydroxides (Berner, 1971a). As nearly all fine-grained terrigenous sediments contain several percent of 'free' ferric iron, the minimum of 1.5% hematite required for red colouration is easily achieved. Hence, all claystones and silty claystones are potentially red. Since, however, most fossil marine claystones are in fact non-red, the content of hematite {or earlier ferric-iron-bearing phases) must have been reduced below the critical level by sedimentary or diagenetic processes. A first important factor is the dilution of pigment by non-red particles. If quartz-silt or carbonate are admixed to a certain a m o u n t of pigment, the hematite content may be taken below the 1.5%-level. This provides one reason for the general absence of red colours from our marine sandstones, or pure carbonates (see also Paul and Franke, 1977). The content of 'free' ferric iron (later hematite) will be highest in the most fine-grained fraction of the sediment, i.e. that fraction which is deposited farther from the source of terrigenous material. This is indicated by the high ferric iron levels and the long stratigraphic range of red shales in these southern parts of the Schiefergebirge which were b e y o n d the reach of coarser clastic sedimentation. The most important factor in the destruction of red pigment is seen in the reduction of ferric iron, which is brought a b o u t as a consequence of oxidation of organic matter. This is clearly d o c u m e n t e d b y reduction spots around conodonts. In the case of reduction bands around silt layers, we suggest that detrital organic matter has been brought in together with the siliceous debris. In addition, the relatively high permeability of silt- and sandstones would have allowed selective impregnation with reducing pore fluids, and thus have enhanced the formation of reduced bands. In reduction spots and bands, the reduced iron has subsequently been removed. This is quite plausible, since most of the organic matter is destroyed prior to compaction, and dissolved c o m p o u n d s could thus easily escape from the system. As is indicated by colour boundaries which cut through the bedding along some distance, and partly follow tectonic patterns, the reduction of iron still continued throughout later diagenesis and into the time of tectonic deformation. This might be due to reactions with remnant organic c o m p o u n d s or pore fluids.

249

The iron reduced at these later stages is, to a varying extent, retained within the system and is incorporated into recrystallized clay minerals or in the dolomitic and ankeritic carbonates, which are to be found in so many Devonian sediments. It is only along open joints that the reduced iron is totally removed. Evidently, the main factors governing the behaviour of the reduced iron are the permeability (intergranular as well as along grosser discontinuities in the rock) and the simultaneous formation (or recrystallisation) of iron-incorporating minerals. Green colouration around calcareous nodules and layers indicates some relationship between the reduction of iron and the precipitation of carbonate. With oxygen present, the decomposition of organic matter yields, besides water, carbon dioxide and this lowers the pH. Carbonate may be dissolved. After the oxygen has been consumed, H2S and ammonia are produced, and the pH is increased (Berner, 1971b). Along with the reduction of iron, carbonate m a y be precipitated, so that reduction spots and bands often occur around concretionary carbonate. Once triggered, the precipitation of carbonate m a y continue, from the surrounding pore fluid, and possibly also from upwelling pore water expelled by compaction. Growing carbonate nodules thus may enclose the front of reduction (that is: the colour boundary), which does n o t migrate further when the decomposition of organic matter has come to an end. Manganese, which is at first reduced and mobilized like the iron, is then preferentially incorporated into the calcite lattice, as the radius of the Mn 2÷ion is closer to that of Ca 2÷ than the radius of Fe ~÷. Thus, the final distribution of manganese does not reflect the result of reduction and solution processes, b u t is controlled b y the precipitation of carbonate. As pointed o u t above, we explain the reduction in spots and bands by the presence of organic compounds. Presumably, this is also true for the thicker non-red portions of the Famennian shale succession, which may have had a higher overall c o n t e n t of organic matter than the intervening red packets. Possibly, the ferric iron which might have gone to the now-green units was already reduced on the shelf, or on the basin floor, b u t the frequent discordance of bedding and colour b o u n d a r y clearly demonstrates that a strong diagenetic overprinting was also possible. Transferring the problem to a larger scale, there remains the question w h y red pelagic mudstones in the Variscan belt are restricted (with few exceptions) to the Upper Devonian and that dominantly in the uppermost Nehden and lower Hemberg Stages. Evidently, it becomes necessary to consider longterm changes in the balance of (1) oxygen supply (in b o t t o m and pore waters) and (2) in the a m o u n t of organic matter admixed in the sediment. In most Devonian and Carboniferous e n v i r o n m e n t s - some euxinic episodes excepted -- it can immediately be observed that there has been enough oxygen in the b o t t o m water to permit some degree of benthonic activity, as is indicated by the occurence of Chondrites-type trace fossils even in dark, basinal shales. Hence, it seems more appropriate to discuss changes in organic

250

supply to the sediment, and we have then to ask why organic productivity in the late Devonian seems to have been significantly lower than during time before and after. The paucity of the fauna, in a well-aerated, normal marine environment, strongly points toward oligotrophy. Some possible reasons, discussed below, necessarily involve a measure of speculation. We assume a minimum of phytoplanktonic productivity during the Famennian, which brought about a minimum of general productivity, and so finally resulted in a decrease in organic supply to the sediment. During the Devonian, the bulk of clastic material as well as of nutrients (such as phosphate) was derived from the Caledonian chains. By late Devonian time, the Caledonian relief was greatly reduced by erosion and the extent of source areas diminished by advancing marine transgressions. This might have caused a deficit of phosphate supply to the open sea, and a subsequent decrease of organic productivity. Similar processes have been envisaged by Tappan (1968, 1970) and by Tappan and Loeblich (1970). The deficit could have been increased by the retention of nutrients in land plants (which conquered the terrestrial regions during the late Devonian) and in subsequently formed soils. A further sink of nutrients might have been due to entrapment of sediment in aerated near-shore and shelf areas, which were enlarged b y the Upper Devonian transgressions. It would also be helpful to our case to assume an o u t p u t of nutrient-rich b o t t o m water and correlated influx of oligotrophic surface water, which would still increase the general oligotrophy, b u t this idea is not substantiated b y any known palaeogeographic configuration. When nutrient supply to the open sea and productivity in the sea were reduced to a minimum, the main source of organic matter was that provided by clastic sedimentation. This is in good agreement with the general coincidence of clastic intercalations and reducing conditions, these last indicated by the drab colours. The combined input of clastic and organic matter was significantly increased in the early Carboniferous, with the onset of the Variscan orogeny, of which one result was the widespread deposition of black shale and greywacke (flysch). It should, however, be said that two other cases of pelagic redbeds are from Cambrian and late Cretaceous flysch sequences (Lajoie and Chagnon, 1973; Faupl and Sauer, 1978), which should likewise be expected to reflect orogenic activity with, in consequence, sufficient nutrient supply from the continent: the deposition of drab or even black shales would have been favoured, if one thinks in terms of the Carboniferous of Germany as a model. This apparent contradiction might be explained by climatic differences. COMPARISONS AND CONCLUSIONS

Critical discussion of the model of redbed development sketched above is somewhat hindered b y the fact that nearly all previous papers deal with

251

terrestrial, fluviatile, or deltaic environments, in which red colouration is often due to breakdown and oxidation of iron-bearing minerals (van Houten, 1973; McBride, 1974; Walker, 1977), b y interstratal solution. In fact, these are the processes that yielded the ferric iron which was finally transported into the Famennian basin, b u t they do not serve to explain the diagenesis in the pelagic environment itself. All models invoke changes and gradients of Eh as the means of producing changes in colour. The importance of permeability in the reduction of pigment has been already noted by several authors (e.g., Lajoie and Chagnon, 1973, p. 99; Thompson, 1950; Horowitz, 1971). The enrichment of red pigment in the clay fraction and decolouration by admixture of silt and sand (dilution) are mentioned (though not discussed in detail) in reviews by Dunbar and Rodgers (1957) and Van H o u t e n (1973). Diagenetic fixation of manganese into calcite has also been observed in pelagic environments elsewhere (Buggisch, 1972; Bencini and Turi, 1974; Faupl and Sauer, 1978). These aspects apart, there are only very few comparable cases in the literature. Lajoie and Chagnon's Cambrian red muds were deposited in a pelagic (flysch) basin with a background sedimentation of green mud. The red detritus is derived from an oxidizing, paralic environment, and is directly and rapidly transported into the basin. This is in contrast to the German Famennian, where the bulk of the clastic material is first deposited on a broad shelf, where iron is reduced after burial, so that the sediments of the outer shelf are non-red and turbidites derived from the shelf edge temporarily establish reducing micro-environments in basins where otherwise the mud would have a red colour. In the Rhenohercynian sea, it is only the clay-size suspended matter that comes directly to the basin floor without experiencing any intervening stage of deposition and reduction. The ferric iron staining Silurian marine redbeds described b y Ziegler and McKerrow (1975) was made available by marine transgressions. Reworked oxidized material was rapidly deposited in marine environments, and remained red because the supply of organic matter was too low to compensate for the strong input of ferric pigment. However, as the Famennian redbeds demonstrate, the red pigment may survive even at low rates of deposition if input of organic matter at the time is also low. The sediments of the Alpine Gosau-flysch (Upper Cretaceous/lower Tertiary) are strikingly similar to the Famennian sequence. Sandstone turbidites are intercalated among pelagic red muds, and green colouration occurs, in a variety of modes, near the coarser layers. Faupl and Sauer (1978) have interpreted the varying positions of the colour boundary in terms of a relationship with changes in the position of the lower limit of oxygen supply, which m a y lie within the b o t t o m water, or at various depths within the sediment. If one takes any of the cases ('Farbabfolgen') in Faupl and Sauer's fig. 3, and repeats the figured members to produce a rhythmic sequence, one arrives at exactly the same picture as for the Famennian: The coarse, basal members of turbidites have a green reduction band, and it is only the vertical

252 reach of the reduction potential that determines the degree of development of the various 'Farbabfolgen'. Hence, we presume that the differential colouration of the Gosau-flysch is entirely of diagenetic origin, is not synsedimentary, and exactly matches the Famennian case in Germany. In a paper on Devonian pelagic red shales in the Harz Mountains, Reichstein (1955) described and correctly interpreted the occurrence of reduction spots and of reduction bands around cherty or silty layers. However, Reichstein excluded terrigenous derivation of the pigmenting iron and favoured volcanic sources instead. Einsele (1963) produced the first comprehensive study of the Famennian basinal clastics, b u t did n o t offer definite conclusions on the origin of the red shales. He assumed that the red pigment is directly imported from the Old Red Continent, and that non-red intercalations have undergone intermittent deposition and reduction on the shelf before they were finally transported into the basin. In our opinion, this is true only of the coarser part of turbiditic suspensions. The drab colouration of clays and silty clays is due instead to diagenetic r e d u c t i o n - a possible process largely neglected b y Einsele. In fact, it is the supply of organic matter and the subsequent processes of diagenetic reduction that cause the decolouration of terrigenous mudstones, all of w h i c h - bearing in mind their original content of ferric i r o n - were potential redbeds. The red pigment may survice in oligotrophic environments, and these occur in various palaeogeographic situations: terrestrial, fluviatile, deltaic, hyperhaline or nearshore marine as well as pelagic-bathyal.

Laboratory procedures The determination of constituents was carried out as follows: total iron, manganese, and phosphate were determined from a HF/HC104 solution. total iron: colorimetry, using ortho-phenantroline (Zimmermann, 1967). ferrous iron: titration with potassium dichromate (Maxwell, 1968}. ferric iron: as numerical difference between total and ferrous iron. manganese: atomic absorption spectroscopy. phosphate: colorimetry, using the vanadate-molybdate method (Koch and Koch, 1964). carbonate: volumetric determinaton of CO2 (Mi~ller and Gastner, 1971). organic carbon: StrShlein-analyzer (coulometric determination of CO2 after combustion, Herrmann and Knake, 1973). extractable iron and manganese: threefold extraction with sodium-dithionite (Scheffer et al., 1961). quantitative mineralogical composition: infrared spectroscopy (Flehmig and Kurze, 1973).

253 ACKNOWLEDGEMENTS

We gratefully acknowledge the help and advice from many colleagues. Thanks for critical reading go to Prof. Dr. D. Meischner, Dr. B. SchulzDobrick, Prof. Dr. K.-H. Wedepohl (all GSttingen), Prof. Dr. H. Ftichtbauer (Bochum), and to Prof. Dr. C. Matthews (Bristol), who also vetted the English manuscript. The great volume of stratigraphic data upon which our palaeogeographic survey is based has been compiled during some decades by a great number of colleagues, whose publications are not mentioned in the text. Dr. B. Schulz-Dobrick carried out the determination of Corg and total CO2 in a number of samples, and Dr. W. Flehmig enabled the quantitative determination of the rock-forming minerals by infrared-spectroscopy. Part of the chemical analyses were carried out by H. Peters. H. Grimme provided the drawings. U. Schmidt-Bucherer typed the manuscript. Funds have been provided by the German Research Society (DFG) through the SFB 48.

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256 Walker, T.R., 1976. Diagenetic origin of continental red beds. In: H. Faike (Editor), The Continental Permian in Central, West, and South Europe. D. Reidel, Dordrecht, pp. 240--282. Ziegler, W., 1971. Conodont stratigraphy of the European Devonian. Mem. Geol. Soc. Am., 127: 227--284. Ziegler, A.M. and McKerrow, W.S., 1975. Silurian marine red beds. Am. J. Sci., 275: 31-56. Zimmermann, M., 1967. Photometrische Metall- und Wasseranalysen. Wissenschaftliche Verlagsgesellschaft, Stuttgart, 3rd. ed.